The American Association of Petroleum Geologists Bulletin

V 64, No. 8 (August 1980) P. 1179-1209, 18 Figs.

Anoxic Environments and Oii Source Bed Genesis'

G. J. DEMAISON^ and G. T. MOORE3

Abstract The anoxic aquatic environment is a mass of water so depleted in oxygen that virtually all aerobic biologic activity has ceased. Anoxic conditions occur where the demand for oxygen in the water column ex­ceeds the supply. Oxygen demand relates to surface biologic productivity, whereas oxygen supply largely depends on water circulation, which is governed by global climatic patterns and the Coriolis force.

Organic matter in sediments ttelow anoxic water is commonly more abundant and more lipid-rich than un­der oxygenated water mainly because of the absence of benthonic scavenging. The specific cause for pref­erential lipid enrichment probably relates to the bio­chemistry of anaerobic bacterial activity. Geochemical-sedimentologic evidence suggests that potential oil source beds are and have been deposited in the geo­logic past in four main anoxic settings as follows.

1. Large anoxic lakes—Permanent stratification pro­motes development of anoxic bottom water, particu­larly in large lakes which are not subject to seasonal overturn, such as Lake Tanganyika. Warm equable cli­matic conditions favor lacustrine anoxia and nonma-rine oil source bed deposition. Conversely, lakes in temperate climates tend to be well oxygenated.

2. Anoxic silled basins—Only those landlocked silled basins with positive water balance tend to become anoxic. Typical are the Baltic and Black Seas. In arid-region seas (Red and Mediterranean Seas), evapora­tion exceeds river inflow, causing negative water bal­ance and well-oxygenated bottom waters. Anoxic con­ditions in silled basins on oceanic shelves also depend upon overall climatic and water-circulation patterns. Silled basins should be prone to oil source bed deposi­tion at times of woridwide transgression, both at high and low paleolatltudes. Sllled-basin geometry, how­ever, does not automatically imply the presence of oil source beds.

3. Anoxic layers caused by upwelling—^These devel­op only when the oxygen supply in deep water cannot match demand owing to high surface biologic produc­tivity. Examples are the Benguela Current and Peru coastal upwelling. No systematic correlation exists be­tween upwelling and anoxic conditions because deep oxygen supply is often sufficient to match strongest demand. Oil source beds and phosphorites resulting from upwelling are present preferentially at low paleo­latltudes and at times of woridwide transgression.

4. Open-ocean anoxic layers—^These are present in the oxygen-minimum layers of the northeastern Pacific and northern Indian Oceans, far from deep, oxygenat­ed polar water sources. They are analogous, on a re­duced scale, to woridwide "oceanic anoxic events" which occurred at ^lot}al climatic warmups and major transgressions^ as in Late Jurassic and middle Creta­ceous times. Known marine oil source bed systems are not randomly distributed in time but tend to coincide with periods of worldwide transgression and oceanic anoxia.

Geochemistry, assisted by paleogeography, can greatly help petroleum exploration by identifying pa-leoanoxic events and therefore widespread oil source bed systems in the stratigraphic record. Recognition of the proposed anoxic models in ancient sedimentary biasins should help In regional stratigraphic mapping of oil shale and oil source beds.

INTRODUCnON The most significant progress made in petro­

leum geology in the last 10 years has been the attainment of a satisfactory understanding of the processes of oil and gas generation and destruc­tion in sedimentary basins. Geochemical tech­niques now routinely identify oil source beds by analysis of rock samples retrieved from deep wells. However, geologists often question the va­lidity of projecting the presence of source beds identified from a few wells and a few grams of rocks, to entire sedimentary basins. Until the last few years these hesitations were legitimate; it was difficult, if not inconceivable, to map oil source beds without a clear understanding of their depo-sitional environment.

Fortunately, recent oceanographic and geo­chemical observations now make it possible to comprehend many formerly obscure aspects of the genesis and stratigraphic distribution of pe­troleum source beds. It is timely, thus, to review depositional environments of oil source beds con­sidering these recent findings and address our­selves to the following questions.

1. What modem sediments are precursors to oil source beds?

©Copyright 1980. The Atnerican Association of Petroleum Geologists. All rights teserved.

AAPG grants permission for i single photocopy of this article for research purposes. Other photocopying not allowed by the 1978 Copyright Law is prohibited. For more than one photocopy of this article, users should send request, article identification number (see below), and $3.00 per copy to Copyright Clearance Center, Inc., 21 Congress Street, Salem, MA 01970.

'The original version of this paper was read at the 10th International Congress on Sedimentology, Jerusalem, 1978, and published in Organic Geochemistry. Manuscript received, August 24, 1979; accepted, January 17, 1980.

^Chevron Overseas Petroleum Inc., San Francisco, California 94105.

^Chevron Oil Field Research Co., La Habra, California 90631.

The writers are indebted to R. R. Hammes. R. W. Jones, R. A. Lagaay, R. S. Oremland, W. D. Redfield, S. R. Silverman, and J. A. Sutherland for critical reviews of the manuscript.

Published with permission of Chevron Overseas Petroleum Inc. and Chevron Oil Field Research Co.

Article Identification Number O149-1423/80/B0O8-0O02$03.0O/0

1179

1180 G. J. Demaison and G. T. Moore

2. What factors affect the preservation of or­ganic matter in aquatic environments?

3. Why are anoxic conditions more favorable than oxic conditions for organic-matter preserva­tion, both in quaUty and quantity?

4. What causes favor anoxic conditions in lakes, seas, and oceans?

5. Can natural anoxic settings be scientifically classified to help oceanographers, stratigraphers, and petroleum explorationists?

PRECURSOR SEDIMENTS TO OIL SOURCE BEDS

Potential oil source beds are organic-rich sedi­ments containing a kerogen type that is sufficient­ly hydrogen-rich (type I or type II; Tissot et al, 1974) to convert mainly to oil during thermal ma­turation. Kerogen type and thus oil source char­acter in ancient sediments is identified through such approaches as elemental analysis of kerogen and whole-rock pyrolysis with additional support from microscopic organic analysis (Tissot and Welte, 1978, p. 81-91; Hunt, 1979, p. 454-472).

Evaluation of organic content, which can be a gross quantitative index of oil generative poten­tial // kerogen is oil-prone, is measured by the amount of organic carbon present in sediments. Documented oil source beds and oil shales around the world always contain hydrogen-rich kerogen and fall into a range of organic carbon content between about 1% and over 20% by weight. The boundary between very rich oil source beds and oil shales is determined by min­ing and processing economics.

Rich to very rich marine oil source beds and oil shales commonly contain higher than average uranium, copper, and nickel concentrations. Ura­nium content in most marine oil shales commonly shows a positive correlation with oil yield upon rock pyrolysis (Swanson, 1960).

The measurement of organic carbon in sedi­ments alone is insufficient to identify potential oil source beds. Transported terrestrial organic mat­ter, oxidized aquatic organic matter, and re­worked organic matter from a previous sedimen­tary cycle can create levels of organic carbon in marine sediments up to about 4%. Yet this abnor­mally concentrated organic matter is hydrogen-poor, gas-prone, and without significant oil gener­ating potential (Tissot et al, 1974; Demaison and Shibaoka, 1975; Dow, 1977). This is essentially the organic facies that has been described in mid­dle Cretaceous marine black shales encountered by several Deep Sea Drilling Project holes in the northwestern Atlantic Basin (Tissot et al, 1979). An identical situation is present today on the Arctic Shelf of the USSR where high organic car­bon concentrations in marine sediments result

from the influx of large amounts of terrestrial or­ganic matter brought in by fluvial discharge (Bezrukov et al, 1977).

In summary, a high organic carbon content in sediments is not necessarily an indication of oil source rock precursor character. Additional geo-chemical evidence, such as measurement of hy­drogen richness of humic substances and kero­gen, pyrolysis yield of whole sediment, and overall soluble-Upid content of the sediment, is needed to ascertain a possible oil source precur­sor character.

FACTORS INFLUENCING ORGANIC MATTER ACCUMULATION IN AQUATIC ENVIRONMENTS

Factors capable of influencing organic-matter accumulation in sediments are both biologic and physical. Biologic factors include primary biolog­ic productivity of the surface-water layers and of adjoining landmasses and biochemical degrada­tion of dead organic matter by metazoan and mi­crobial scavengers. Physical factors include modes of transit of organic matter to depositional sites, sediment particle size, and sedimentation rates. These factors interact to determine quaUta-tive and quantitative preservation of organic mat­ter in sediments.

Primary Biologic Productivity The principal source of aquatic organic matter

is the phytoplankton (Bordovsky, 1965) com­posed largely of single-cell microscopic algae re­siding in the uppermost layers of water illuminat­ed by sunlight, the euphotic zone. The main limiting factor to planktonic productivity, in ad­dition to light, is the availability of mineral nu­trients, particularly nitrates and phosphates, which are in short supply in the euphotic zone. Phytoplankton are intensively grazed by zoo-plankton. Both phytoplankton and zooplankton are then consumed by large invertebrates and fish.

The other source of organic matter in the aquatic environment is transported terrestrial or­ganic matter from streams and rivers. Land-plant productivity is largely dependent on the amount of rainfall on supporting landmasses. Because ter­restrial organic matter has undergone consider­able degradation in subaerial soils prior to its transport, it is usually hydrogen depleted and re­fractory in nature.

The traditional view is that fields of high sur­face productivity in the ocean should be associat­ed with high organic enrichment of underlying sediments; however, after exhaustive investiga­tion, we could not find a systematic correlation between primary biologic productivity (Fig. 1)

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Anoxic Environments 1183

and organic carbon content of bottom sediments in the oceans (Fig. 2). Some areas of high produc­tivity like the Peruvian Shelf and Southwest Shelf of Africa were enriched; others of equally high productivity like the Grand Banks of Newfound­land, the northeastern Brazilian Shelf, or the Northwest Shelf of Africa were not. Clearly, fac­tors other than surface productivity had to be evaluated.

Biochemical Degradatioii of Dead Organic IVfatter

Dead organic matter is inherently unstable thermodynamically and seeks its lowest level of free energy in any given environment. Above all, it serves as a source of energy and nutrients for Uving organisms. Heterotrophic microorganisms, mainly bacteria, play a critical role in decompos­ing organic matter in the water column, in the interstitial water of sediments, and in the di­gestive tracts of animal scavengers. In marine ba­sins bacterial biomass approaches that of the phy-toplankton (Bordovsky, 1965). Bacterial degradation proceeds quickly and efficiently in aerobic (oxygen-rich) water. The overall oxida­tion processes by aerobic bacteria are illustrated by:

(CH2O) + O2 ^ CO2 H- H2O.

When the oxygen supply becomes exhausted, oxidation of organic matter by anaerobic bacteria using nitrates as a source of oxygen (electron ac­ceptor) takes over according to the generalized scheme:

(CH2O) -I- 4NO3 -^ 6C02-I- 6H2O -I- 2N2.

Once nitrates have been exhausted, degrada­tion of organic matter continues by anaerobic bacteria using sulfates as the oxidant by the gen­eralized reaction scheme:

CH2O -I- SO4 ->• CO2 -I- H2O -I- H2S.

The last and least efficient step in anaerobic metabohsm is fermentation. Here carboxyl (CO2) groups and organic acids of the organic matter itself, or from bacterial breakdown, are employed as electron acceptors. A special type of anaerobic fermentation is bacterial methanogenesis (Clay-pool and Kaplan, 1974). It is a complex process carried out in several steps by types of bacteria which degrade one another's by-products (Man-heim, 1976). The methane-generating (methano-genic) bacteria occupy the terminal niche of the anaerobic food web and produce methane by CO2 reduction by H2 or by attacking acetate, for­mate, or methanol (Wolfe, 1971). Higher hydro­

carbons, at least up to butane, are always present together with bacterial methane in anoxic water, although in much smaller amounts (Deuser, 1975). Anaerobic methane fermentation is com­mon below the sulfate reducing zone in recent marine sediments and prevails in anoxic waters where both nitrate and sulfate levels are low (e.g., in freshwater lakes). Methane fermentation and sulfate reduction, however, are not mutually ex­clusive (Oremland and Taylor, 1978). The pro­cesses mentioned are presented in a simplified form. They represent an important developing area of research in geomicrobiology.

Anaerobic degradation is thermodynamically less efficient than aerobic decomposition (Clay-pool and Kaplan, 1974) and results in a more lip-id-rich and more reduced (hydrogen-rich) organic residue than does aerobic degradation (Force and McCarty, 1970; Beliaeva and Romankevich, 1976; Gerschanovich et al, 1976; Pelet and De-byser, 1977; Didyk et al, 1978). Moreover, under such conditions, a significant fraction of the pre­served organic matter consists of remains of the bacterial biomass itself (Lijmbach, 1975). The mechanisms leading to this preferential enrich­ment in lipids are Uttle understood and in need of fundamental research. The enhanced preserva­tion of hydrogen-rich and lipid-rich organic mat­ter in sediments deposited under anoxic water is critical for the genesis of oil source beds.

Explanations proposed for the higher than nor­mal organic matter concentrations observed un­der anoxic water (Richards, 1970; Deuser, 1971) include the suggestions that anaerobic degrada­tion is inherently slower (Lijmbach, 1975) or that enhanced preservation might be a consequence of high sedimentation rates (Richards, 1970). (The latter is discussed under the effects of other physi­cal parameters.)

As far as biologic factors are concerned, Force and McCarty (1970) have shown by laboratory experiments that rates of anaerobic degradation of algae by sulfate reduction are identical to those of aerobic degradation. Observations in the natu­ral environment by Orr and Gaines (1974) led to parallel conclusions. This suggests that most of the organic matter consumption at oxygen-rich ("oxic") sediment/water interfaces is by the meta-zoan, not by the microbial population (Degens and Mopper, 1976).

Under an oxic water column (Fig. 3) the benth-ic fauna actively scavenges and reworks the "or­ganic rain" falling through the water column. Ex­cept in shallow waters where sunlight penetrates to the bottom, there is very Uttle primary produc­tion and mainly consumption of organic matter at the benthic boundary. Bottom muds under an oxygen-rich water column are commonly anoxic

1184 G. J. Demaison and G. T. Moore

OXIC ENVIRONMENT

TYPE OF RESPIRATION

RESIDENCE TIME

OF O.M.

POORER CM. PRESERVATION (0.2-4%T.O.C.) LOWER QUALITY O.M.

BIOLOGICAL REWORKING IS ENHANCED BY:

. PRESENCE OF ANIMAL SCAVENGERS AT INTERFACE.

. BIOTURBATION BY WORMS FACILITATES DIFFUSION OF OXIDANTS (02,S04) IN SEDIMENTS.

. LESSER ORGANIC COMPLEXATION WITH TOXIC METALS.

OXYGEN CONSUMPTION

days-mos.

SULFATE 1 REDUCTION ••

BACTERIAL CO2 REDUCTION

750 yrs

500 yrs 6

750 yrs ^

U=50cm/1000yrs

FIG. 3—Degradation of organic matter under oxygenated (oxic) water column. Organic carbon concentration under oxygen-rich water is closely related to sedimentation rate, without regard to surface primary productivity.

but nevertheless extensively disrupted by oxygen-respiring invertebrates such as polychaetes (worms), holothurians, and bivalves. The most common modes of feeding are on particulate or­ganic matter and bacteria in water above the sedi­ment (suspension feeders), and on organic matter and bacteria present within the sediment itself (deposit feeders). Mobile deposit feeders (burrow-ers) cause mixing and transport of particles, as well as irrigation of sediments, thereby accelerat­ing geochemical processes at or below the sedi­ment-water interface (Rhoads, 1974; Aller, 1978). Many who have studied deep-sea sediments be­lieve that the sediments are mixed to a depth of 5 to 30 cm by biologic activity (Peng et al, 1977). Bioturbation is ubiquitous under oxic water where it has been observed at all water depths, including the deep-sea sediments of the abyssal realm. The range of mixing (bioturbation) rates in the deep sea is comparable to that in nearshore regions, with no apparent correlation between the mixing rate and the sediment type or sedimenta­tion rate (Turekian et al, 1978). This suggests that mixing rate may be almost solely related to benthic biomass variabihty, and thus to produc­tivity in surface waters and subsequent dehvery of edible organic material to the ocean floor.

Under an anoxic water column (Fig. 4), how­

ever, oxygen depletion (even in the absence of hydrogen sulfide) depresses and eventually eUmi-nates benthonic metazoan Ufe (Theede et al, 1969). The benthic metazoan biomass is unaffect­ed by oxygen concentrations as low as about 1 ml/1 (Rhoads, 1974), but between 0.7 ml/1 and 0.3 ml/l it sharply decreases by about a factor of 5 (Rosenberg, 1977). Below 0.3 ml/1, deposit feed­ers become rare, less active, and soft bodied only. Eventually bioturbation ceases. Below 0.1 ml/l, the suspension feeders disappear, leaving anaero­bic bacteria as the only effective reworkers of or­ganic matter.

Once organic matter is incorporated in the anoxic sediment itself, the lack of bioturbation acts as a limiting factor to diffusion of oxidants into the sediment—hence bacterial sulfate reduc­tion is slowed down if not completely arrested. A classic observation is that the pore fluids of anox­ic sediments are depleted in sulfates (Manhetm, 1976, p. 115-118). Lack of bioturbation under anoxic water results in laminated and organic-rich sediments. Reducing conditions imder anox­ic water make certain toxic metals, like lead, available for chelation with organic matter (Jones, 1973). Degens and Mopper (1976) have suggested that metal complexation of organic matter makes it less susceptible to microbial at-

Anoxic Environments 1185

ANOXIC ENVIRONMENT BETTER O.M. PRESERVATION {1-25%T.0.C.)

HIGHER QUALITY O.M.

BIOLOGICAL REWORKING IS SLOWED BY:

.THE ABSENCE OF ANIMAL SCAVENGERS.

. RESTRICTED DIFFUSION OF OXIDANTS (SO4) INTO UNDISTURBED SEDIMENT.

.LESSER UTILIZATION OF LIPIDS BY ANAEROBIC BACTERIA.?

TYPE OF RESPIRATION

OXYGEN CONSUMPTION

RESIDENCE TIME

OF O.M.

BACTERIAL SULFATE

REDUCTION

days - mos.

OCEAN BOTTOM

BACTERIAL CO, REDUCTION

•100 cm •

U=50cm/lOOOyrs

FIG. 4—Degradation of organic matter under anoxic water column. Bioturbation becomes minimal at oxygen concentrations below 0.5 ml/1. This concentration, rather than total absence of oxygen, is effective "biochemical fence" between poor and good organic matter preservation.

tack. Comparison of Figures 3 and 4 shows that, giv­

en identical sedimentation rates, dead organic matter in oxygenated environments is exposed to oxidants much longer than in anoxic environ­ments, owing to bioturbation. For a given sedi­mentation rate, bioturbation imder oxic water prolongs the exposure of organic matter to oxida­tion by hundreds of years, which strongly en­hances degradation. Added to this effect is the grazing of the sediment bacterial biomass by mo­bile deposit feeders responsible for bioturbation.

IVfo«|es of Tnmsit of Organic Matter to Depositknial Sites

E>ead zooplankton, unconsumed algal cells, fe­cal pellets, and fish carcasses originating in the shallow euphotic zone continuously sink to the bottom. The speed of fall of particulate organic matter is exceedingly slow and ranges from 0.10 to 5 m per day, according to their shape and size, the smallest particles being the slowest to reach bottom. The transit time for fecal pellets is much faster (Spencer et al, 1978). It is estimated that their passage through a water column 4 km deep will not exceed 15 days.

If organic debris falls through a well-oxygenat­

ed water column in which animal life is abun­dant, it is consumed by scavenging animals, until much of the final organic "rain" which reaches the bottom consists of fecal pellets (Spencer et al, 1978). These are reworked biologically at the sed­iment interface, thereby losing their original mor­phologies.

Long residence of organic matter in the water colimm before sedimentation adversely affects its preservation (Degens and Mopper, 1976). Thus the depth of the water column as well as the size of organic particles affects the quantity and qual­ity of the organic input to the sediments.

Influence ai Sediment-Particle Size In addition to being affected by bioturbation,

bacterial activity in recent sediments is also influ­enced by granulometry. Fine-grained sediments, where diffusibility of oxidizing agents is re­strained, have lower levels of bacterial activity than coarse-grained sediments (Bordovsky, 1965), which helps explain the broad correlation be­tween particle size and organic carbon content in modem and ancient sediments. Besides being de­posited in high-energy environments, coarse clas­tic sediments, such as sands, permit easy diffusion of free oxygen and oxidizing salts dissolved in wa-

1186 G. J. Demaison and G. T. Moore

ter (Tissot and Welte, 1978, p. 81-91). They are always very low in organic matter.

Influence of Sedimentation Rate The rate of accumulation of organic carbon in

marine sediments is closely related to the bulk accumulation rate (Heath et al, 1977). This regu­larity is well demonstrated under oxic water in areas such as the Argentine basin (Stevenson and Cheng, 1972), the northwest African continental margin (Hartmann et al, 1976), and the deep re­gions of the North Pacific and North Atlantic Oceans (Heath et al, 1977). In all these areas there is a positive correlation between sedimentation rate and organic carbon content in sediments. The latter usually ranges between 0.3 and about 4%. Under anoxic water this correlation is less evident, as will be discussed later. The overall range of organic carbon under anoxic water in modem sediments is significantly wider and high­er (about 1 to 20-1- %), regardless of sedimenta­tion rate.

Furthermore, insufficient documentation exists for modem sediments with regard to possible di­lution effects created by very high sedimentation rates (above 500 cm/lO' year). There is evidence that the positive correlation between organic car­bon content and sedimentation rate at low and intermediate sedimentation rates may not extend indefinitely into the realm of very high rates. If it did, prodelta muds should be the richest organic sediments in the marine realm, yet organic car­bon contents are systematically below 1% in mod­em prodelta sediments of the Louisiana Gulf Coast (Dow and Pearson, 1975), Niger delta (Klingebiel, 1976), and Amazon delta (Bezrukov et al, 1977) pointing to a dilution effect. Primary biologic productivities offshore from the Missis­sippi and Amazon deltas are exceptionally high (Fig. 1) but, as expected in such oxic environ­ments, are not reflected by the sedimentary or­ganic carbon content.

Sediments under anoxic water tend to be or­ganically richer than those under oxic water, largely due to lack of benthonic scavenging and absence of bioturbation at the seafloor. Of even greater importance to the genesis of oil source beds, sedimentary organic matter is more reduced and lipid-rich under anoxic water than under oxic water. The cause for this preferential lipid enrich­ment probably relates to the biochemistry of anaerobic bacterial activity and requires further research.

Under oxygenated ("oxic") water, fluctuations in organic carbon content are clearly related to sedimentation rates (up to the point where dilu­tion becomes significant), with little influence of surface productivity. Variations in surface pro­

ductivity appear systematically, compensated by quantitative variations of the benthic biomass re­sponsible for organic matter consumption and bioturbation at the seafloor. Under anoxic water, both sedimentation rates and perhaps surface productivity are factors which explain fluctuation in organic carbon content within the observed range. The respective inputs of these two factors, however, are yet to be satisfactorily elucidated.

DEVELOPMENT OF ANOXIC CONDITIONS IN WATER MASSES

Maximum oxygen saturation in sea water is about 6 to 8.5 ml/1, depending on water tempera­ture and salinity. For this study we define as "anoxic" any water containing less than 0.5 millil­iters of oxygen per liter of water (0.5 ml/1), which is the threshold below which the metazoan benth­ic biomass and, more specifically, bioturbation by deposit feeders, becomes significantly depressed. Therefore, it is proposed as the effective "bio­chemical fence" between potentially poor or good quaUtative and quantitative preservation of or­ganic matter in sediments.

There are two "end member" causes to anoxia in natural waters: excessive oxygen demand and deficient oxygen supply. Anoxic conditions occur where the natural demand for oxygen in water exceeds the supply.

Oxygen Demand in Natural Water Oxygen consumption in water is essentially a

biochemical process resulting from the degrada­tion of organic matter produced in the shallow layers of the euphotic zone. At least 80% of this oxidation occurs in these shallow layers and de­creases sharply with depth. However, dead organ­ic matter that has escaped total degradation and has sunk to the bottom continues to create a de­mand for oxygen which, however weak in relation to that prevailing in the euphotic zone, remains to be satisfied by a matching supply. If this contin­ued demand is not replenished by deep-water cir­culation, the water colunm becomes anoxic.

Even water columns with a normal oxygen sup­ply can become anoxic. This occurs in areas of very high primary productivity wherever the oxy­gen supply near sea bottom cannot cope with the load of descending dead organic matter.

Oxygen Supply in Natural Waters

Oxygen is supplied to water masses by two physical processes: (1) downward movement of oxygen-saturated water from the well-aerated sur­face layers (oxygen is supplied to the surface lay­ers by exchange with the atmosphere and by pho-tosynthetic oxygen production) and (2) upward movement of oxygen-rich, colder, denser bottom

Anoxic Environments 1187

water into intermediate water zones. Three physical properties of water govern oxy­

genation of bottom waters: water density increas­es with increasing salinity; water density increas­es with decreasing temperature (to 4°C); and oxygen solubility in water varies inversely with decreasing temperature and salinity.

If oxygen-rich surface water becomes denser because it is saltier or cooler than the surrounding water, it sinks to the bottom and circulates there as an aerating undercurrent and causes multilay-ered vertical stratification in water bodies of all sizes. The terms, thermocUne, pycnocline, and ha-locline, describe temperature, density, and salini­ty boundaries, respectively.

Stratification of seas and oceans in terms of oxygen concentration has long been recognized by oceanographers. Modem oceans would be en­tirely anoxic at depth without aeration of their basins by colder and denser oxygen-rich bottom water derived from the pwlar regions, mainly from the high southern latitudes.

The circulation of these deep oxygen-rich bot­tom waters is only partly known. Oxygen supply to the intermediate and deep oceanic water de­pends on patterns of water circulation due to sur­face-wind stress, density differences, high-latitude cooling, and the Coriolis force (caused by the earth's rotation). The dynamics of general ocean­ic circulation are complex, fluctuating, and not yet fully understood quantitatively.

The most common cause of anoxia is the inca­pacity of the oxygen supply in water to meet the biochemical oxygen demand. Hence lack of verti­cal mixing and oxygen renewal in deep waters is, perhaps, the most important factor controlling the location of anoxic layers and thus, indirectly, the preferred sites of deposition of oil source bed precursors. In the words of Wyrtki (1962), "Bio­chemical processes are responsible for the exis­tence of oxygen minima, but circulation is re­sponsible for the position."

CLASSIFICATION OF ANOXIC ENVIRONMENTS The world map of organic carbon (Fig. 2),

based mainly on Bezrukov et al (1977), does not distinguish the areas with organic carbon enrich­ment over 3%. Thus input of transported terrestri­al organic matter as well as the effects of high sedimentation rates under oxic water cannot be separated from enrichment in marine organic matter created by anoxic conditions. In the writ­ers' experience with ancient marine sediments, or­ganic carbon contents rarely exceed 3% when transported humic material of terrestrial origin is predominant. Conversely, ancient marine sedi­ments containing more than 3% organic carbon always contain a significant portion of aquatic or­

ganic matter. Consequentiy, we reviewed in detail studies of

recent sediments reported by previous researchers to investigate the implications of enrichments higher than 3%, potentially caused by anoxic con­ditions. We found it possible to classify present-day anoxic environments into four main types: (1) large anoxic lakes; (2) anoxic silled basins; (3) anoxic layers caused by upwelhng; and (4) the anoxic open ocean.

LARGE ANOXIC LAKES Oxygen depletion in inland seas and large lakes

is determined by the supply-demand balance be­tween free oxygen availability in bottom water and planktonic productivity of organic matter in the shallow layers.

Plant nutrients such as phosphates and nitrates are carried into lakes and inland seas by fluvial drainage systems that transport solutes leached from soils. These nutrients usually hmit the planktonic productivity of lakes, which then de­termines the amount of oxygen needed to recycle dead organic matter. Eutrophic lakes are charac­terized by an abundance of dissolved plant nu­trients and by a seasonal oxygen deficiency in the bottom waters. Oligotrophic lakes are deficient in plant nutrients and contain abundant dissolved oxygen in their bottom water.

Oxygen supply in the bottom waters is usually good in areas of contrasting climate with seasonal overturning of the lake (Swain, 1970, p. 73-111). Also, cold, well-oxygenated stream and river wa­ter sinks to the bottom and enhances oxic condi­tions. Oxygen supply is lower in warm tropical climates because of lack of seasonal overturn and lower oxygen contents due to higher water tem­peratures.

Examples of Large Anoxic Lakes The two best studied anoxic lakes in the world

are Lake Kivu and Lake Tanganyika, both part of the East African rift-lake system.

Lake Tanganyika (Degens et al, 1971)—Lake Tanganyika, proposed as the type example for "large anoxic lakes" (Fig. 5), measures 650 km by up to 70 km. The maximum water depth is about 1,500 m and anoxic conditions lethal to metazoan hfe prevail below about 150 m. Some hydrogen sulfide is present in the anoxic water. Sediment cores from both deep and shallow water show considerable vertical variabiUty or varving.

Sediments deposited in the shallow aerated wa­ter within the lake contain 1 to 2% organic car­bon, whereas the deep-water anoxic sediments range between 7 and 11% in organic carbon con­tent. Carbon isotope ratios of recently deposited organic matter in anoxic sediments of Lake Tan-

1188 G. J. Demaison and G. T. Moore

LAKE TANGANYIKA NW SE

CH4,H^S PERMANENT THERMOCLINE

LONGITUDINAL SECTION 525 Km

Km

YJ5 1 15

ANOXIC SEP. ORG. CARB. 7 -11 % OXIC SEP. ORG. CARB. 1-2%

U = 5 - 50 cm /lOOOyrs

SIMILAR SETTINGS:

LAKE KIVU and DEAD SEA

FIG. 5—Type example of "large anoxic lake,' sedimentation rates in cm/1,000 years.

Lake Tanganyika. Symbol U in this and following figures expresses

ganyika are in a range normally associated with organic matter deposited in marine sediments: - 2 1 to - 2 2 %o (Pedee belemnite—PDB). Re­mains of diatoms are the most abundant fossils in these sediments.

Lake Kivu (Degens et al, 1973}—-Lake Kivu is nearly 500 m deep but is entirely anoxic below a depth of about 60 m. The pH of the surface water is around 9 but drops below 7 in the anoxic deep water where hydrogen sulfide is present in detect­able quantities. The recent deep-water sediments under the anoxic zone contain up to 15% organic carbon.

The distribution of organic carbon in associa­tion with the stratigraphy of carbonate and anhy­drite beds, in the past 6,000 years, reveals signifi­cant cUmatic fluctuations. Sapropel-rich beds coincide with humid climate and high water. Pe­riods of dry climate, low water, and evaporites coincide with lower, but still significant, organic enrichments.

Examples of Large Oxic Lakes

Some East African rift lakes are oxic, notable examples being Lake Albert and Lake Victoria. Lake Albert (Tailing, 1963) is a thermally strati­fied lake 50 m deep. There are a few oxygen-de­pleted pockets down to 0.5 ml/1. The shallow depth and density currents of cooler, oxygenated water are responsible for good ventilation. Lake Victoria is large but shallow (about 100 m) and is affected by seasonal overturns. It is not anoxic (Swain, 1970).

All the very large temperate lakes of the north-em hemisphere are well oxygenated. Three exam­ples are typical: (1) Lake Baikal, in Siberia, al­though the world's deepest lake (1,620 m), is well aerated and no anoxic conditions are present (Swain, 1970); (2) the Great Lakes of North America are oxic (Swain, 1970), although Lake Erie may be partly and intermittently anoxic be­cause of excessive nutrient supply by pollution (phosphates from detergents and nitrates from sewage water); and (3) Great Slave Lake and other large lakes in western Canada are also well oxygenated (Swain, 1970).

Anoxic Enriroiuiients in Large Lake Systems

Tropical climate, with little seasonal change and moderate to high rainfall year around, pro­motes permanent water column stratification and therefore anoxia, particularly if water depth is in excess of 100 m. The sediments deposited are commonly both carbonate- and organic-rich. Or­ganic carbon content is higher (commonly over 10%) than in anoxic marine sediments which may be caused by anaerobic degradation in fresh wa­ter being largely by fermentation rather than by the more efficient sulfate reduction process pre­vailing in anoxic marine environments.

Large anoxic lakes are not present in cold and temperate climates. Seasonal overturning of wa­ter, the higher capacity of cold water to dissolve oxygen, and density underflows of cold river wa­ters all enhance oxic conditions in bottom lake waters of temperate and cold regions.

Anoxic Environments 1189

ETtdence of Large Anoxic Lakes fai Past Geologk Time A spectacular example of past oil shale and oil

source bed deposition in large anoxic lake sys­tems is the Eocene Green River Formation in the western interior United States (mainly Colorado and Utah).

The Green River "oil shales" are brown, highly laminated, dolomitic marlstones containing hy­drogen-rich kerogen (type I; Tissot et al, 1974). Retorting and conversion of this kerogen, during heat treatment of the whole rock, yields commer­cial shale oil. The Mahogany zone comprises the richest oil shales of the Green River Formation.

Several chemical and sedimentologic character­istics of these Mahogany zone "oil shales" dis­cussed by Smith and Robb (1973) clearly demon­strate that permanent anoxic conditions existed at lake bottom during deposition. Smith and Robb's conclusions can be summarized as follows.

1. Evidence for complete lack of bioturbation is given by the absence of benthic fossils and the presence of minute seasonal laminations (varves) that can be followed laterally for tens of miles. Such features imply persistent lack of water movement and a lethal environment for benthic fauna in the bottom water.

2. The mineral composition of the "oil shale" implies a very alkaline, sodium carbonate-rich, bottom water whose higher density enhanced per­manent chemical stratification of the lake. Eh-pH chemical fences (Krumbein and Garrels, 1952) implied by Mahogany "oil shale" mineral and chemical composition point to high pH (alkaline) and highly reducing (anoxic) conditions in the bottom water. These were lethal to macrohfe forms, thus explaining the lack of bioturbation and fine varving of the "oil shales." The upper surface layer of the lake above the density bound­ary was fully oxygenated and supported plank-tonic life, for organic matter was continually de­posited in the sediment.

The lower part of the Green River Formation, below the Mahogany "oil shales," contains zones of finely laminated or varved, papery, kerogenous shales which usually assay less than 15 gallons per ton of rock (Roehler, 1974). These older anoxic shales differ from the Mahogany "oil shales" in that they were deposited in a freshwater lacus­trine environment instead of an alkaline lake.

The lower part of the Green River Formation, mainly below the Mahogany zone "oil shales," has reached the principal stage of oil generation and is responsible for most of the crude oils pro­duced in the Uinta basin (Tissot et al, 1978).

Crude oils generated from anoxic lake beds such as those present in the Green River Forma­tion are highly paraffinic and typically feature

high pour points, when undergraded, and very low sulfur contents. The nonmarine Lower Creta­ceous oils of China, Brazil, and some coastal west African basins belong in this category.

ANOXIC SILLED BASINS All silled basins have several physical barriers

(or sills) that restrict vertical mixing, thereby en­hancing water stratification. Silling alone is not sufficient to create anoxic bottom conditions; certain patterns of water circulation, largely con­trolled by climate, need also to be present (Gras-shoff, 1975).

Basins with a positive water balance have a strong salinity contrast between fresher outflow­ing surface water and deeper ingoing more saline and nutrient-rich oceanic water. The devel­opment of permanently or intermittently anoxic conditions is a general feature of those semi-enclosed seas which have a positive water bal­ance. Basins with a positive water balance also act as nutrient traps, thus enhancing both produc­tivity and preservation of organic matter. Typical examples of anoxic landlocked silled basins with a positive water balance in a humid zone are: the Black Sea (Degens and Ross, 1974); the Baltic Sea (Grosshoff, 1975); Lake Maracaibo (Red-field, 1958); and Saanich Inlet, British Columbia (Nissenbaum et al, 1972).

In basins with a negative water balance, result­ing from a hot, arid climate, there is a constant inflow of shallow oceanic water to compensate for high levels of evaporation. As shallow oceanic water enters, it replaces the hypersaline water which sinks and flows out as a density undercur­rent into the ocean. Therefore, the basin bottom is both oxygenated and nutrient depleted (Fig. 6). Examples of oxic, negative-water-balance, silled basins without significant bottom organic enrich­ment are the Red Sea, Mediterranean Sea, and Persian Gulf.

Black Sea The Black Sea is the largest anoxic landlocked

basin in the world. It has been studied in greater detail by oceanographers and geochemists than any other anoxic basin (Degens and Ross, 1974; Usher and Supko, 1978). We propose the Black Sea as the type example of an anoxic silled basin.

The Black Sea has a positive water balance with an excess outflow of fresh water resulting in relatively low salinity of surface water. As a re­sult, a permanent halocline is present which also marks the boundary between oxic and anoxic conditions. Hydrogen sulfide is present in the anoxic zone. Its upper boundary is slightly con­vex, lying at a depth of about 250 m around the edges of the Black Sea and rising to about 150 m

1190 G. J. Demaison and G. T. Moore

EVAPORATION

OXYGENATED SALTIER DENSER WATER SINKS

DENSER WATER SPILLS OUT

LEAST OXYGENATED WATER

OXIC "NEGATIVE WATER BALANCE" BASIN EXAMPLES: MEDITERRANEAN - RED SEA - PERSIAN GULF

FIG. 6-Schematic model of oxic "negative water balance" basin. Hypersaline water flows out at depth over sill because it is denser than open oceanic water. This underflowing current keeps the basin bottom well ventilated and nutrient depleted.

in the center (Fig. 7). Below those depths the en­vironment is lethal to all fish and invertebrate life.

About 22,000 years ago the Black Sea was a freshwater lake. Approximately 11,000 years ago rising Mediterranean waters began to invade the Black Sea as a consequence of climatic warm-up and ice retreat. About 7,000 years ago, the hydro­gen sulfide zone began to form and present con­ditions were established about 3,000 years ago.

When anoxic conditions began about 7,000 years ago, the maximum organic carbon content of the sediments changed from 0.7 to 20% (Hirst, 1974). The organic-rich microlaminated black layer (7,000 to 3,000 years ago) is about 40 cm thick and is called "the old Black Sea" or "unit 2" (Degens and Ross, 1974). Unit 2 sediments com­monly exceed 1% organic carbon.

Unit 1, between 3,000 years ago and the pres­ent, is about 30 cm thick and is composed of al­ternating white coccolith- and black organic-rich microlaminates (50 to 100 laminae per centime­ter). Organic carbon content in unit 1 is still sig­nificant, but lower (about 1 to 6%) than in the underlying unit 2. In the two deepest basins, or­ganic carbon contents in unit 1 range between 2 and over 5% (Fig. 8).

On the basis of varve chronology (Degens et al, 1978), rates of sedimentation in the deep basin were 10 cm/W years for the sapropel-rich unit 2 and 30 cm/lO' years for the more recent cocco-

lith-rich unit 1. Thus the organically richer sedi­ments (unit 2) actually correspond to lower sedi­mentation rates. The correlation is inverse to that found in deep-sea sediments under oxygenated water by Heath et al (1977).

A sedimentation rate map for the whole Black Sea during the last 3,000 years (Ross and Degens, 1974) shows a possible overall negative correla­tion between sedimentation rate and organic mat­ter concentrations (Shimkus and Trimonis, 1974). Sedimentation rates on the organic-lean upper slope of the Black Sea are higher by a factor of three than in the organically rich deep basins.

Further, Shimkus and Trimonis (1974) ob­served that "there is no correlation between phy-toplankton production and organic-matter con­tent in sediments" of the Black Sea. "Fields of high organic-matter content correspond to areas of low primary production and vice-versa."

Recent geochemical investigations by Pelet and Debyser (1977) on a suite of Black Sea oxic unit 3 and anoxic unit 2 sediments showed that: (1) in both units organic matter is mixed, about % ter­restrial plants and % marine plankton; and (2) the anoxic sediments (unit 2) have organic carbon contents from 3.85 to 14.95%, whereas the oxic sediments (unit 3) contain 0.65 to 0.69% organic carbon. The humic acid plus fulvic acid contents in relation to organic carbon are twice as high in oxic unit 3 than in anoxic unit 2. Hydrogen/car-

Anoxic Environments 1191

BLACK SEA A "POSITIVE WATER BALANCE BASIN" N

BOSPORUS

SILl? (27m.)

PERMANENT HALOC! INF f20q

H.'S.CH^

Km 0 .5

M ^J5

2 5 0 Km

ANOXIC SEP. ORG. GARB. 1 - 1 5 % OXIC SEP. ORG. GARB. <2 .5%

U: 5 - 3 0 c m / 1 0 0 0 Y r s

SIMILAR SETTINGS: BALTIC SEA • SAANICH INLET

LAKE MARACAIBO

FIG. 7—Type example of "anoxic silled basin," Black Sea. When anoxic conditions began, about 7,000 years ago, average organic carbon content in sediments increased about tenfold. Lipid content in relation to organic carbon also significantly increased in anoxic sediments.

FIG. 8—Content of organic carbon in modem sediments of Black Sea (from Shimkus and Trimonis, 1974). Highest organic carbon concentrations occur in deep abyssal plains. Apparently no correlation is between organic carbon concentration (this figure) and sedimentation rate (Ross and Degens, 1974) in recent anoxic sediments of Black Sea.

1192 G. J. Demaison and G. T. Moore

bon ratios of humic compounds are higher in sed­iments from anoxic unit 2 than in the oxic sedi­ments of underlying unit 3. Significantly, the chloroform extractable lipid content, in relation to organic carbon, is five times higher in sedi­ments from the anoxic unit 2 than in the oxic sediments of unit 3.

The organic carbon budget of the Black Sea has been estimated by Deuser (1971). He con­cluded that about 80% of the organic input is re­cycled in the top 200 m of water. The remainder falls into the anoxic hydrogen sulfide-poisoned water where approximately half of this organic material is further degraded by sulfate-reducing bacteria, which leaves about 5% of the original input for incorporation into the sediment and 5% solubilized in the anoxic water. Therefore, even in the most favorable model for organic-matter pres­ervation (under anoxic water), 95% of the organic matter escapes fossilization and is eventually re­cycled.

BaMcSea

The Baltic Sea, the largest brackish water area in the world, has a positive water balance (Gras-shoff, 1975). It is affected by a permanent halo-cline with a pronounced dip to the east, from the Kattegat into the Gulf of Bothnia. A permanent oxygen deficiency exists below the halocline, with intermittent anoxia and hydrogen sulfide poison­ing being present in the lower part of the water column in the Gothland Deep. Increase in bot­tom-water oxygen depletion during the last 75 years may be due partly to man-made pollution. Organic carbon maps are available for the Baltic Sea (Romankevich, 1977) and show a pattern of organic carbon enrichment above 3% coincident with the areas where the water column is anoxic.

Examples (rf Oxic Silled Basins

Two of the world's largest silled basins, the Red Sea and the Mediterranean Sea, are well oxygen­ated at depth and their modem sediments tend to be organic-poor. Both these basins are character­ized by a negative water balance, that is, a larger inflow from the ocean than output to the ocean. The negative balance results from loss due to evaporation (Fig. 6).

In the Red Sea, oxygen-rich shallow water flows in from the Gulf of Aden and moves all the way up to the Gulf of Suez (Grasshoff, 1975). As it becomes saltier and denser, the water sinks and returns back at depth into the Gulf of Aden over the shallow sill of Bab el Mandeb. Localized oxy­gen depletion in the southern Red Sea is devel­oped at the foot of the sill just before the dense, deep bottom water spills over southward into the Gulf of Aden. The northern end of the Red Sea,

farthest away from the sill, shows the highest lev­els of oxygenation in bottom water. Organic mat­ter is low throughout the Red Sea, and sediments are largely biogenic clastic carbonates.

The Mediterranean is the world's largest land­locked silled marine basin, but bottom water is now entirely oxygenated at depth (Fairbridge, 1966). The lack of anoxic conditions in the Medi­terranean Sea is due to a negative water balance circulation pattern identical to that observed in the Red Sea. There is, however, evidence from cores taken by DSDP and other oceanographic expeditions of past anoxic events in the eastern Mediterranean during the Pleistocene (Stanley, 1978). Intermittent anoxic conditions prevailed five times during the last 9,000 years. They are possibly the result of large and sudden influxes of fresh water, perhaps due to an increase in precipi­tations and/or ice retreat, which intermittently turned the eastern Mediterranean into a Black Sea-like, positive water balance basin.

Compared to oxidized Pleistocene sediments which contain less than 0.5% organic carbon, the anoxic muds of the upper Pleistocene of the east-em Mediterranean contain over 10 times more organic carbon (2 to 8%; Fairbridge, 1966). The latter are also laminated and devoid of benthic fauna.

The Persian Gulf is another silled basin with a negative water balance because of arid climate and intense evaporation. Bottom waters are oxy­gen-rich because of deep hypersaline water out­flow and the highest organic carbon values mea­sured (up to 2.5%) are in the silty clays that floor the deepest depressions in the southern part of the Persian Gulf (Hartmaim et al, 1971).

Baffin Bay, between Canada and Greenland, is a large silled marine basin where bottom water exhibits mild oxygen depletion (about 3 ml/1) but is nowhere near anoxic conditions. Dense, cold, oxygen-rich arctic water from the northern sill (Nares Strait) slowly sweeps the bottom in pulses and eventually spills out of Davis Strait, the southem sill, and into the Labrador Sea (Palfrey and Day, 1968).

Silled Depressioiis in Oxic Open Ocean

The world's oceans contain many silled depres­sions. They are well oxygenated except in rare examples like the Cariaco Trench and the Orca basin.

Cariaco Trench—The Cariaco Trench is a clas­sic example of a relatively small local anoxic de­pression located on an oxygenated open ocean shelf (Richards and Vaccaro, 1956). This small silled depression, with a maximum depth of 1,400 m, is located on the continental margin north of Venezuela. It is anoxic below 300 m, the sill depth

Anoxic Environments 1193

being about 200 m. On the open oxygenated shelf, organic richness of recent sediments aver­ages about 0.8% organic carbon. At the bottom of the trench, where anoxic conditions prevail, or­ganic carbon concentrations increase as much as eight times in relation to the oxic environment, with values reaching 6% (Gormly and Sackett, 1977). Benthic foraminifers are lacking in the laminated anoxic sediments, but fish bones are very abundant. Radiocarbon measurements and studies of pelagic foraminifers indicate that stag­nation of the trench was synchronous with an abrupt warming of surface water about 11,000 years ago (Richards, 1976).

Detailed geochemical investigations of the Car-iaco Trench anoxic sediments have been reported by Combaz and Pelet (1978) together with identi­cal studies on sediments deposited under oxic wa­ter on the Demerara Abyssal Plain and on the outer deltaic cone of the Amazon River.

The anoxic sediments of the Cariaco Trench (1) are organically richer than their sedimentation rate (50 to 75 cm/10' years; Combaz and Pelet, 1978) would permit under oxic water; (2) contain insoluble organic matter that is hydrogen-rich, with kerogen in the "oil-prone" type II, category (the Amazon and Demerara kerogens and humic acids are less rich in hydrogen than their Cariaco Trench equivalents; they fit in the type III "gas-prone" category); and (3) contain larger amounts of hydrocarbons and other lipids, in relation to total organic carbon, than the oxic Amazon and Demerara sediments.

Lastly, investigation of fatty-acid content in the interstitial waters of the Cariaco, Demerara, and Amazon sediments, as compared to overlying sea-water (Saliot, 1977), shows preferential enrich­ment in fatty acids of the anoxic sediments in the Cariaco Trench.

Orca basin—The Orca basin is on the lower continental slope in the northwest Gulf of Mexico where complex bathymetry is largely attributed to salt diapirism and slumping of sediments. The ba­sin is a crescent-shaped depression enclosing an area of approximately 400 sq km. The sill depth is approximately 1,900 m on the southeastern flank. The basin has two deeps slightly greater than 2,400 m at the north and south ends of the cres­cent.

The water in the Orca basin below a depth of 2,200 m, or about 200 m above the basin floor, is a brine. In addition to the pronounced increase in salinity at 2,200 m, there is a reversal of the tem­perature gradient, and an abrupt decUne in oxy­gen, from 5 ml/1 from a sample of 100 m above the halocline, to 0.0 ml/1 below the halocline.

Oxygen, as well as nitrate, is depleted by bac­terial activity (Shokes et al, 1977). In all probabil­

ity the brine in the Orca basin came from solution of a salt diapir at or near the surface somewhere around the basin slo(>e but the precise location of this hypothetical feature is not known.

Organic carbon in cored sediments under the anoxic water ranges between 0.8 and 2.9% (Sack­ett et al, 1978). Some of these values may be af­fected by dilution from slumped-in sediments from the fairly steep sides of the basin (Sackett et al, 1978).

EvMence of Anoxic Silled Basins in Past Geologic Time

Many anoxic silled basins are suspected in the geologic record but few are well documented in terms of integrated environmental and geochemi­cal studies. Thus, in Great Britain, study of sedi-mentologic and fossil features in the Lower Juras­sic (Toarcian) of Yorkshire by Morris (1979) permitted him to subdivide an apparently monot­onous shale sequence into three facies (Fig. 9)— normal, restricted, and bituminous.

The normal shale is a homogeneous, bioturbat-ed sediment with abundant benthic body-fossils and common sideritic nodules. This facies is in­dicative of well-oxygenated bottom waters. The restricted shale consists of poorly laminated sedi­ments with scattered calcareous concretions, sparse benthic fauna, and thin discrete pyritic burrows; this facies was deposited under oxygen-depleted water, perhaps close to 0.5 ml/1. The bi­tuminous shale is a finely laminated sediment with pyritic concretions, little or no bioturbation, and a benthic fauna which is sparse and does not include burrowing organisms; this facies reflects anoxic conditions in the water column.

The three facies described by Morris (1979) oc­cur in cycles and are indicative of variations in the position of the oxic-anoxic boundary within the water column (Fig. 9), replicating the tyjje set­tings described for modem sediments in Figures 3 and 4. We applied geochemical kerogen typing by Rock-Eval pyrolysis (Espitalie et al, 1977; Clem-entz et al, 1979) combined with organic carbon measurements on samples representative of the three facies, kindly supplied by the University of Reading (United Kingdom). The results are sum­marized on Figure 9; they confirm that the "nor­mal shale" contains only "gas-prone" type III kerogen whereas the restricted shale is a mixed organic facies (type Ill-type II). Only the anoxic "bituminous" shale contains highly "oil-prone" type II kerogen. Widespread laminated oil shales in the Lower Jurassic (Toarcian) of the Paris ba­sin as documented by Hue (1978) also resulted from anoxic basin conditions.

Alternating coccolith limestones, bioturbated clays, and laminated organic-rich oil shales de-

1194 G. J. Demaison and G. T. Moore

SHALE FACIES

AND KEROGEN TYPE

GEOCHEMICAL DATA

BIVALVE

GROUPS

TRACE

FOSSILS CONCRETIONS

ENVIRONMENTAL

INTERPRETATION

NORMAL SHALE

T Y P E I I I "GAS-PRONE"

KEROGEN

C ORG %:

0 6 6 - 3 . 4 0

HYDROGEN INDEX:

83-134

EPIFAUNAL a

I N F A U N A L

SUSPENSION,

INFAUNAL

DEPOSIT

ABUNDANT

CHONDRITES

HORIZONTAL

BURROWS

SIDERITIC

a CALCAREOUS

Oxic BoHom Woter

Oxidizing Conditions

teducing *" ' * —

%,« cing Conditions <sVOs<

RESTRICTED SHALE

TYPE n -nr " M I X E D "

K E R O G E N

C0RG7o'

2 .59-6 .75

HYDROGEN INDEX:

135-216

DOMINANT

INFAUNAL

DEPOSIT

FEW

UNBRANCHED

HORIZONTAL

BURROWS

CALCAREOUS W e a k l y O x i c Bot torn WQ t e r

BITUMINOUS S H A L E

TYPE n "OILPRONE"

K E R O G E N

CORG%:

5.61-11.42

HYDROGEN INDEX:

2 5 3 - 5 8 4

DOMINANT

EPIFAUNAL

SUSPENSION

NONE

NO BURROWS

PYRITIC

CALCAREOUS Anoxic Bottom Water.

FIG. 9—Interpretation of facies of Liassic (lower Toarcian) of Yorkshire, Great Britain. This format, excepting geocbemical data, was adapted from Morris (1979). Hydrogen Index, ratio expressing hydrogen richness of kerogen (Espitalie et al, 1977). Investigated rocks are all within same 60-m (197 ft) thick surface section. Degree of matura­tion of organic matter, as determined from Rock-Eval pyrolysis temperatures, is late immature to early mature.

posited during Late Jurassic time in the Kimmer-idgian of southern England (Dorset) show a strong analogy with recent Black Sea sediments. This association reflects cycUc oscillations of an oxic-anoxic boundary within a stratified water column (Tyson et al, 1979) during Late Jurassic time. Kinuneridgian to Volgian organic-rich anoxic shales are also present under parts of the North Sea. Where thermally mature, because of sufficient burial, they are the major source contri­butors to the North Sea giant oil reserves (Ziegler, 1979).

Anoxic, bituminous shales of very widespread areal extent were also deposited at approximately the same time (Volgian rather than Kimmeridgi-an) in the Western Siberian basin (Kontorovich, 1971). These Upper Jurassic anoxic shales, where thermally mature, are the source of most of the oil reserves entrapped in the many giant fields of this prolific petroleum province.

The best documented example of a former anoxic silled basin in North America is the Mow-ry sea which occupied in latest Albian time most of the northwestern interior United States. Byers and Larson (1979) described three distinct sedi-mentologic facies in the Mo wry Shale: (1) lami­nated mudstone, (2) bioturbated mudstone, and (3) bioturbated sandstone. The three facies repre­sent the following environments: (1) a low-ener­

gy, lethal environment, under anoxic water, where bioturbation is absent; (2) a low-energy en­vironment where bioturbation is present because of sufficient oxygenation of the water column; and (3) a high-energy, intensely bioturbated and well-oxygenated environment. These were de­scribed by Byers and Larson under the Schaefer (1972) terminology of marine biotopes as (1) le­thal isostrate, (2) vital isostrate, and (3) vital hete-rostrate.

The organically richest zones identified by Nix­on (1973) as oil source beds of major significance coincide with the low-energy, unbioturbated, le­thal facies, which was deposited under anoxic wa­ter in the Mowry sea.

ANOXIC LAYERS CAUSED BY UPWELLING

Upwelling is a process of vertical water motion in the sea wherever subsurface water rises toward the surface. Large-scale upwelling due to wind-stress-induced, Ekman transport occurs along certain coastUnes such as those of Cahfornia, Peru, Chile, South-West Africa, Morocco, and Western Australia. Upwelled water in coastal re­gions comes from relatively shallow depth, usual­ly less than 200 m.

Ziegler et al (1979) described three distinct geo­graphic situations where favorable conditions are met for coastal upwelUng: (1) meridional upwell-

Anoxic Environments 1195

S.W. AFRICAN SHELF

%o02

ANOXIC PATCHES EXTEND ABOUT 700 Km ALONG SHELF

Km 0 .5

• 1

•1.5 2

ANOXIC SEP. ORG.CARB. 3-26% OXIC SED. ORG. CARB. < 3 %

SilMILAR SETTiNG: PERU

FIG. 10—Type example of "anoxic layers caused by upwelling," South-West African Shelf (Benguela Current). Coastal upwellings are complex current and counter-current systems interacting with offshore winds. Upwelled water, rich in nutrients and low in oxygen, does not originate from great water depths, but from less than 200 m.

ing on north-south coasts between 10 and 40° lat. on the east side of oceans (e.g., California Cur­rent); (2) zonal upweUing on east-west coasts of equator-centered continents at about 15° lat. in association with the easterly trade winds (e.g., Caribbean Ciurent); and (3) monsoonal upwell­ings on diagonal east-facing coasts of equator-centered continents at about 15° lat. in associa­tion with easterly trade winds (e.g., Somalia Cur­rent).

Upwellings can also occur in the open ocean at water-mass boundaries, and such have been not­ed around the Antarctic continent, along the equator, and between Iceland and Norway. Up­welled water is rich in nutrients (nitrates and phosphates) and, therefore, promotes high biolog­ic productivity. Recycling of dead organic matter in the water column creates a very high oxygen demand which can trigger anoxic conditions in deeper water layers under the upwelling. A classic example of anoxicity in the water column (essen­tially created by excessive biochemical oxygen demand due to coastal upwelling) occurs on the shelf offshore South-West Africa (Namibia) in as­sociation with the Benguela Current. We propose it as our type example (Fig. 10).

Benguela Cuirent

Considerable oceanographic and geochemical research has been conducted offshore of South-West Africa since the mid-1960s, notably by Cal­vert and Price (1971a, b). Anoxic conditions oc­

cur on the South-West African Shelf (Fig. 10), particularly off Walvis Bay. The oxygen-depleted zone at sea bottom is an elongate area approxi­mately 50 by 340 km (Calvert and Price, 1971a) parallel with and close to the coastline. Beyond the shelf break (about 100 km offshore) normal oxygenated conditions return to the bottom, as the regional oxygen-minimum layer in the Atlan­tic, south of Walvis Ridge, is only very weakly developed (Bubnov, 1966). Anoxic conditions are created on this narrow shelf by the high oxygen demand from decomposition of large amoimts of plankton resulting from the Benguela Current up­welling. The upwelling is due to a combination of a cold coastal current (the Benguela Current) and persistent offshore winds blowing northwest. Shallow surface water is skimmed off by the wind, permitting nutrient-rich subsurface water to ascend from a depth of about 200 m.

In three dimensions the system visualized is one where oxygen-poor, but nutrient-rich water constantly moves up and mixes in the euphotic zone with oxygenated water, causing high biolog­ic productivity along a narrow coastal band. Dead plankton eventually falls to the bottom un­der the upwelling, and nutrients associated with the organic matter are brought back to the sur­face by the upweUing instead of being dispersed into the open sea. WTiat is buried and lost to the bottom sediments is replaced by nutrients brought in by the Benguela Current.

Brongersma-Sanders (1972) wrote: "Upwell-

1196 G. J. Demaison and G. T. Moore

ORGANIC CARBON CVJ

• /SYLVIA HILL

FIG. 11—Content of organic carbon in modem sedi­ments offshore Walvis Bay (Namibia; from Calvert and Price, 1971). Highest values are off Walvis Bay and Cape Cross, where anoxic conditions and bacterial ni­trate reduction intermittently prevail in bottom waters.

ings are fertile in the first place because nutrient-rich water is brought to the well-lighted surface layers, but this is not the only or even the main point, which is that upwelling is a kind of counter current system. A counter current system acts as a trap in which nutrients tend to accumulate."

Organic carbon concentrations in sediments under the oxygen-depleted zone range between 5 and 24% (Fig. 11). The highest values (over 20%) are off Walvis Bay and Cape Cross where oxy­gen-depleted conditions reach the anoxic level

(below 0.5 ml/1). Bacterial nitrate reduction oc­curs in bottom water. Free hydrogen sulfide has been detected in the water of Walvis Bay, causing occasional mass mortahty of fish and detectable hydrogen sulfide odor inland. There appears to be a clear positive correlation between level of oxygen depletion in the bottom water and organic carbon enrichment of the sediment.

Most organic matter is planktonic in origin be­cause runoff from the land is intermittent and from a practically rainless desert (Calvert and Price, 1971b). Sedimentation rates were calculat­ed by Veeh et al (1974) for anoxic sediments off­shore Walvis Bay. Rates range between 29 and 103 cm/lO^ years. Organic carbon values range between 5.65 and 16.36% in the same samples and show no systematic correlation with sedimenta­tion rates.

The anoxic, organic-rich sediments in the Ben-guela Current contain abnormal concentrations of copper, nickel, uranium, and phosphorus. Cop­per, molybdenum, and nickel distributions show similarities to that of organic carbon (Calvert and Price, 1971a). Uranium concentrations in sedi­ments of the South-West African Shelf have been investigated by Veeh et al (1974). They reached the following conclusions: (1) the uranium has a positive correlation with organic carbon; (2) the uranium was derived from modem seawater; and (3) fixation of uranium in the sediments was con­ditioned by the presence of phosphorus together with reducing (anoxic) conditions; uranium is in­corporated into apatite and is concentrated in phosphatic nodules and laminae.

When upwellings occur in areas of broad and well-developed intermediate oxygen-depleted lay­ers, the oxygen starvation brou^t about by the local upwelling reinforces the regional oxygen anomaly. Notable coastal upwellings of this type have been investigated off California and Peru in association with regionally oxygen-depleted inter­mediate water layers.

Peru Current The Peru (Humboldt) Current is a well-docu­

mented example of an anoxic environment asso­ciated with coastal upwelling (Fig. 12; Fairbridge, 1966). The Peru Current refers to a system of rela­tively shallow currents flowing northward along the west coast of South America. It is a complex system involving two surface currents, an under­current, and a countercurrent (Idyll, 1973).

The degree of upweUing offshore of Peru is re­lated to the intensity of the wind stress. Prevailing trade winds off the coasts of northern Chile, Peru, and Ecuador blow principally from the northeast and south. This wind flow pushes the surface wa-

w

Anoxic Environments

PERUVIAN SHELF

1197

100 Km.

ANOXIC BAND ~XTENDS ALONG

HELP FOR OVER 1000 Km.

ANOXIC SEDL ORaCARR 3-11% OXIC SEP. ORG. CARa 05-3%

Km. rO .5

•1

•2

•3

L5

FIG. 12—Example of "anoxic layers caused by upwelling," Peru Current. Geochemical observations on sediments at bottom of Peru Shelf and Trench clearly demonstrate preferential preservation of lipids in those sediments under anoxic water.

ter northward at the same time that the CorioUs force deflects it to the west, thus skimming off the surface layers and letting cold subsurface water well up. Study of the isotherm patterns estabhshes that the depth of upwelling ranges from about 50 to 240 m (Fairbridge, 1966).

The upwelled water is undersaturated in oxy­gen but rich in nutrients (phosphates, nitrates). The ample supply of nutrients is associated with an exceedingly high rate of primary productivity. The combination of high productivity in the sur­face waters and depleted oxygen in the water col­umn generates conditions favorable to enhanced preservation of organic matter in the underlying marine sediments. The concentration is further enhanced by the lack of significant water move­ment toward the ocean. TTie Peru Current, like the Benguela Current, is a self-regenerating nu­trient trap and an outstanding primary producer of organic matter.

The organic matter in the bottom sediments of the Peruvian Shelf has been investigated by Ger-shanovich et al (1976). The highest concentra­tions of organic carbon (average 3.33% with val­ues up to 11%) are in silty clays and laminated diatomaceous oozes under the anoxic zone at wa­ter depths between 100 and 500 m (Figs. 12, 13).

Organic-matter quality of its alkali-extractable portion was measured by elemental analysis. Hy­drogen/carbon ratios of this humic material reach 1.37 to 1.43, indicating that they are not fulvic and humic acids identical to those of soils

or peats but are related to the "sapropelic acids" of Russian researchers. This type of marine hum­ic material is considered as the main precursor of oil-prone kerogens. Hydrogen richness of kerogen precursor material is lower (H/C = 1.23) in the more oxygenated bottom sediments of the trench below 500 m. The same oxygenated sediments contain three times less extractable lipids in rela­tion to organic carbon than those under the anox­ic zone. Concurrently, the ratio of total lipid to total organic carbon of the bottom waters versus the sediments was investigated for the anoxic and oxic zones by Beliaeva and Romankevich (1976). Under oxygenated bottom waters in the trench, the soluble lipid to organic carbon ratio decreases by a factor of 22 from the water column into the sediments. On the anoxic shelf, however, the same ratio only decreases by a factor of 7. These observations clearly demonstrate preferential preservation of lipids in sediments under anoxic water.

The sediments of the Peruvian Shelf deposited under anoxic water contain significantly in­creased quantities of sapropel-type organic mat­ter (70 to 90% of the organic matter); enhanced contents of soluble organic matter (bitumen or total lipids) and significant quantities of hydro­carbon gases including higher hydrocarbons up to hexane (Gershanovich et al, 1976) are also pres­ent. Uranium-enriched phosphorite nodules are present along the lower and upper boundaries of the oxygen-minimum layer at the edges of the

1198 G. J. Demaison and G. T. Moore

FIG. 13—Content of organic carbon in recent sedi­ments offshore Peru (from Logvinenko and Romanke-vich, 1973).

anoxic zone (Veeh et al, 1973).

Examples of Upwellings WHhout Anoxic Lay«s

Not all areas of upwelling and high primary productivity are associated with anoxic layers in the water column and good concomitant organic-matter preservation in bottom sediments. Many areas of high primary productivity are underlain by highly oxygenated water which permits thorough aerobic degradation of the organic mat­ter. Documented examples of such upwelling zones where the oxygen supply at sea bottom ex­ceeds the biochemical oxygen demand are in Ant­arctica, the northern Pacific, and off southeastern Brazil.

In Antarctica, despite locally high surface pro­

ductivity, the circumpolar waters are perhaps the most highly oxygenated on earth, largely because of very low surface temperatures (Fig. 14). Organ­ic carbon values in bottom sediments around Antarctica reported in the Geological-Geophysical Atlas of the Indian Ocean (Udintsev, 1975) range between 0.10 and 0.94%.

In the northern Pacific, offshore Japan and the Kuriles, very high productivity caused by upwell­ing is offset by a constantly renewed supply of oxygen-rich intermediate and deep water coming from the Bering Sea.

Concerning offshore southeastern Brazil, Sum-merhayes et al (1976) concluded after deUiled study, "Beneath the biologically productive up-welled surface waters, there is remarkably httle sedimentation of organic matter. Also, although phosphorite deposits are usually associated with upwelhng centers, there are no phosphatic sedi­ments off southeastern Brazil . . . this probably results from the high degree of oxygenation of upwelled water which contains 6 ml/1 of dis­solved oxygen."

EvMence of Anoxic Layen Caused by Upwelling in Past Geologic Time

Paleogeographic reconstructions as applied to occurrence of Phanerozoic marine phosphorites (Ziegler et al, 1979) and petroleum source beds in the Paleozoic (Parrish et al, 1979) clearly point to the following.

1. Phanerozoic phosphorites were deposited in association with upwelling zones at low paleolati-tudes.

2. Phanerozoic phosphorites are not randomly distributed in time. Most deposits were formed in upwelling zones during periods of worldwide transgression and expansion of the oxygen-mini­mum layer.

3. Phosphorites and organic-rich sediments are commonly associated, as in the Permian Phos-phoria Formation (Parrish et al, 1979) in North America. The Phosphoria black shale members are major source contributors to Paleozoic oil ac­cumulations in the western interior United States (Claypool et al, 1978; Momper and Williams, 1979), including several giant fields.

Evidence of anoxic sedimentation associated with coastal upwelling zones also is present in the Tertiary of California. Sedimentary suites charac­teristic of past deposition under the oxygen mini­mum layer of an upwelling are present at various levels in all the oil producing basins of Cahfomia. These sediments typically fit the following de­scriptions: laminated, phosphatic, bituminous, dark-brown shales, and laminated, hard, sili­ceous, dark-brown, commonly diatomaceous shales. Uranium-bearing phosphatic material is

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1200 G. J. Demaison and G. T. Moore

abundant throughout, whether in nodules or fine­ly disseminated, and radioactive background is relatively high. These rocks are generally devoid of macrofossils.

Anoxic organic-rich, bituminous, phosphatic shales are best developed in: the lower Pliocene (base of the Repetto Formation in the Los Ange­les basin); the upper Miocene (lower part of die Puente Formation "nodular shale" in the Los An­geles basin, also the McLure and Antelope shales in the San Joaquin Valley); and the middle Mio­cene (major parts of the Monterey Formation in the Ventura basin, Santa Maria basin, and south-em Coast Ranges; middle Miocene, i.e., Luisian, time marks the peak of phosphatic deposition in the Tertiary of California; Dickert, 1966). The formations listed have been geologically recog­nized as prolific oil sources in the oil-producing California basins, long before the advent of mod­em petroleum geochemistry (Jenkins, 1943).

The most widespread anoxic sediments caused by past upwelling in California are in the middle Miocene (Luisian Stage) as well as in the lower part of the upper Miocene (Mohnian Stage). These stages, particularly the Mohnian, corre­spond with highest sea stands and maximum transgression during Miocene time on a global scale (Vail and Mitchum, 1979). Anoxic condi­tions associated with upwelling were also rein­forced locally, in some places, by silled-basin to­pography. The overall oceanographic setting was clearly similar, however, to that present today in the Pera Current upwelling.

ANOXIC OPEN OCEAN

The intermediate layers of the open ocean both north and south of the equator in the eastern Pa­cific and northem Indian Ocean have been found to be oxygen depleted over areas of considerable size (Fig. 14). The volumes of anoxic water in today's open oceans greatly exceed the volumes of anoxic water present in the Black Sea.

It has been suggested that the distribution and position of these widespread oxygen-minimum layers result essentially from biochemical oxygen demand created by high planktonic productivity. Oceanographic observations on a global scale, however, do not agree with this hypothesis. Bio­chemical oxygen demand is indeed the triggering mechanism of anoxicity in general, but deep cir­culation patterns govern the distribution and po­sition of these regional anomaUes.

The Grand Banks of Newfoundland in the northwest Atlantic form one of the world's great fishing grounds with productivities comparable to those offshore California or northwest Africa, yet no oxygen depletion of the intermediate or bot­tom water has been observed there. The overall

water column is well oxygenated (over 6 ml/1) and organic carbon in bottom sediments is pres­ent only in low concentrations, averaging 0.5% (Emery and Uchupi, 1972, p. 292, 371-373). The overall lack of correlation between areas of high oceanic productivity and oxygen depletion is also clearly demonstrated in the Pacific and Indian Oceans (Figs. 1, 14). The oxygen-depleted areas in these oceans are cul-de-sacs or are in the lee of sources of cold oxygenated bottom water. Oxygen concentrations decrease very gradually between Antarctica and the northem Indian Ocean and northeast Pacific (Reid, 1%5; Gorshkov, 1974) demonstrating very broad regional circulation ef­fects, rather t^an local anomaUes related to over-productive coastal upwellings.

In only two important regions do substantial amounts of cold oxygenated surface water contin­uously sink to abyssal depth: the North Atlantic and around Antarctica. TTie deep oxygen concen­tration in the other oceans diminishes with in­creasing distance from these polar sources. The deep, cold, oxygen-rich bottom-water underflow from the poles counterbalances the evaporation from shallow waters in the subtropical regions of the oceans.

Strong bottom-current activity is prevalent in the cold, high-latitude regions and on the westem side of basins; these are the areas of most turbid bottom water (HoUister et al, 1978). Red pelagic clays also occur on the westem floors of oceanic basins (Fairbridge, 1966). These observations are coincident with better overall oxygenation of bot­tom and intermediate water at high latitudes and on the western sides of the Pacific and Atlantic Oceans (Fig. 14). This unequal latitudinal distri­bution of the regional oxygen-minimum layer (Fig. 14) indicates a tendency toward stagnation on the eastern sides of oceanic basins. It reflects, according to Stommel (1958), the effects of the Coriolis force upon deep circulation of oxygenat­ed bottom water. It is clearly unrelated to areas of high primary productivity present on both sides of the same oceanic basins. Because this asymmetry is created by the earth's rotation, the same effect must have persisted throughout geologic time. Thus it could have had a profound influence on the paleogeographic distribution of oil source beds on continental shelves and slopes.

Significant organic enrichment has been ob­served around the world wherever the regional oxygen-minimum layer falls at or below 0.5 ml/l and intersects the continental slope and shelf. This correlation has been recognized and de­scribed in the northeastem Pacific Basin (Gross et al, 1972), in the Gulf of California (van Andel, 1964), and in the Indian Ocean (Stackelberg, 1972; Konjukhov, 1976).

Anoxic Environments 1201

INDIAN OCEAN

ANOXIC INTERME­DIATE LAYER (< 0.5)

r - ^ UPWELLING

SIIMILAR SETTING:

N.W. PACIFIC

CROZET BASIN

ARABIAN BASIN

4 mil/liter OXYGEN IN WATER

FIG. 15—Type example of "anoxic open ocean," Indian Ocean. Highest organic carbon concentrations are present wherever oxygen minimum layer faUs at or below 0.5 ml/1 and intersects continental slope and shelf.

Indian Ooean

The Indian Ocean is the type example of the "anoxic open ocean" of this classification.

The upper continental slope of the Indian Ocean from the Gulf of Aden to the Andaman Islands is occupied by a very large and relatively shallow anoxic layer marked in vertical hachures on Figure 15. Wherever this layer impinges on the shelf and slope between 250 and 1,200 m (Fig. 16), abnormally high organic carbon concentra­tions (between 2 and 10%) have been found by Stackelberg (1972), Konjukhov (1976), and other investigators (Udintsev, 1975). Organic carbon content on the shelf and other parts of the slope under oxic water is lower (between 0.5 and 1%). Similar observations were made in the Gulf of Oman (Hartmann et al, 1971).

Primary productivity varies from high, related to upwelling (marked in horizontal hachures on Fig. 15) on the west side of India, to low in the Gulf of Bengal. Sedimentation rates also vary widely from low, off the Arabian coast and in the Gulf of Oman, to very high off the Indus and Ganges deltas (Lisitzin, 1972, p. 135-171). The strongest correlation which overcomes differences in proiductivity and sedimentation rates is that be­tween anoxic conditions and organic enrichment.

Pacific Ocean Offsiiote Waaidngton-Oiegon

In the northeast Pacific, offshore Washington and Oregon, a positive correlation is evident be­tween organic carbon content of bottom sedi­ments and oxygen depletion of bottom water (Gross et al, 1972). Where the dissolved-oxygen concentrations fall below 1 ml/1, organic carbon values range between 1 and 3%. Between 1 and 2 ml/1, organic carbon contents of bottom sedi­ments fall to around 1%. Other factors such as grain size and sedimentation rate are significant, but it is statistically clear that the highest organic carbon concentrations occur where the oxygen-minimum layer impinges on the continental slope. In this general area, "there is surprisingly httle correlation between known areas of high pri­mary productivity at the ocean surface and the amount of organic matter preserved in the sedi­ment" (Gross et al, 1972).

Gulf of CaUfomia

The recent sediments of the Gulf of California (Mexico) have been investigated in detail by Cal­vert (1964), Parker (1964), Phleger (1964), and van Andel (1964) in relation to oxygen content of intermediate and deep waters (Roden, 1964) and primary productivity of surface water. Determi-

1202 G. J. Demaison and G. T. Moore

Water Depth ^2 1 Z 3 It 5mlA_

Mineral sond Relicts ond \ J I Vr^\,

Mica a R. ijiud X x V I C o ' - g . I Corb. jfecpell

tooo

2000

3oqo_

FIG. 16—Section through oxygen minimum layer on western shelf of India (modified from Stackelberg, 1972).

nations of organic carbon were made on approxi­mately 100 core samples throughout this area. The organic carbon content appeared at first to correlate to the amount of clay in the sample. However, inspection of samples with comparable clay contents showed that the largest amounts of carbon (5% and above) are not in the silled basins within the gulf itself, but on the slopes of the southern gulf and in the open sea. In these areas the oxygen-minimum layer in the Pacific Ocean (below 0.5 ml/1) impinges on the slope between about 300 and 1,200 m. These areas of high or­ganic enrichment and anoxic conditions in the water column coincide with finely laminated, richly diatomaceous sediments and the absence or scarcity of burrowing organisms. Conversely, lack of lamination due to homogenization by ac­tive bioturbation, as well as lower organic carbon content (below 2%), were demonstrated for those portions of the slopte where oxygen concentra­tions in the water column are higher than 0.5 ml/1 (Fig. 17).

Calvert (1964) also observed that oxygen con­centrations lower than 0.5 ml/1 in the water col­

umn did not completely inhibit the development of benthonic foraminifers but virtually sup­pressed the existence of macroinvertebrates and burrowing organisms.

Adandc Ocean

Although an oxygen-minimum layer has long been recognized here as well, the Atlantic Ocean in general tends to be oxygenated, for it is an open corridor between the two polar regions where cold, oxygen-rich bottom waters originate. There are only two notable oxygen-depleted in­termediate zones, both in the eastern South At­lantic basins (Fig. 14). They are on the west side of Africa, on both sides of the equator, one off Cape Verde (Senegal) and one in the Angola ba­sin (Bubnov, 1966). Only the Angola basin reach­es the anoxic level (0.5 ml/l), but detailed geo-chemical observations on bottom sediments are lacking.

Offshore northwest Africa (Mauritania and Cape Verde), dissolved-oxygen concentrations in the oxygen-minimum layer are still in the oxic range (above 1 ml/1) whereas surface productivity

Anoxic Environments 1203

200-

400-

600'

^ 800

1000-X >-S I ZOO-

IE

5 1400-

1600'

1800'

2000

NUMBER OF CORES I 2 3 4 5 6 7 8 9 10 II 12 13 14 IS 16 I I I t I 1 I I 1 I I t I t I I

ZJ >.<iH(ni If 1 111 [I

H y H g i l l ' ••'• >f •• '

ifH'' t

OXYGEN CONCENTRATION ml/L 0.5 I 2 3 4 5 6

I I 1-

^ g LAMINATED m SEDIMENTS

f ' i .V'- .V

FIG. 17—Degree of bioturbation in relation to oxygen concentration in bottom water observed in Gulf of Califor­nia (modified from Calvert, 1964). Laminated sediments are present at oxygen concentrations lower than 0.5 ml/1. This level of oxygen depletion, not total absence of oxygen, is effective threshold to lack of bioturbation.

is high owing to local upwelling. Organic carbon measurements by Gaskell et al (1975), Hartmann et al (1976), and Debyser et al (1977) demonstrate that organic preservation is variable in this oxic environment with a range of less than 0.5 to about 4% organic carbon. These organic carbon fluctuations under oxic water are controlled by sedimentation rates and do not reflect present high productivities created by upwelling condi­tions (Hartmann et al, 1976).

Offshore Ivory Coast and Ghana, oxygen de­pletion is also present in the intermediate water (about 1.5 ml/1), but does not reach the anoxic level. Organic carbon values range between 3 and 5% but organic matter is largely composed of transported terrestrial plant remains (Klingebiel, 1976). The predominance of transported terrestri­al organic matter, originating from rain forests, into an open-marine environment under 1,200 m of water, is an interesting present-day example of a situation common in ancient open-marine sedi­ments. Such a setting implies abundant rainfall on nearby continental masses.

In the Niger delta, in nearly identical climatic surroundings, organic carbon content in prodelta clays is low and ranges between 0.5 and 1.0% (Klingebiel, 1976). Sedimentation rates are very

high, suggesting that dilution of organic matter must occur there.

Evidence of Widespread Anoxic Open-Ocean Conditions in Past Geologic Time

Oceanographic observations are made today on a cold earth. Present time represents a rela­tively brief oscillation within the Pleistocene Ice Age sequence. The modern oceans and seas are relatively well oxygenated at depth because con­trasted climatic belts and polar ice caps enhance vigorous convective water mass movements (Berggren and HoUister, 1977). The active bottom circulation that characterizes the present Atlantic, for example, was not initiated until about 3 m.y. ago. Q)ld-earth conditions as we know them to­day were heralded by major cUmatic cooling dur­ing late Eocene-Oligocene time followed by late Miocene and Phocene-Pleistocene global marine regressions and glaciations. The long-term trend since Eocene time is one of global climatic cool­ing concurrent with a lowering of the mean sea level in the world ocean (Vail and Mitchum, 1979). In the geologic past, deep oceanic circula­tion was not so active during periods of warm equable global cUmate which corresponded with high mean sea levels and the absence or extreme

1204 G. J. Demaison and G. T. Moore

reduction of polar ice caps. For example, sedi­mentary and paleontologic data interpreted with­in the framework of plate tectonics suggest that the Arctic Ocean and north-polar regions were ice-free in Late Jurassic (Hallam, 1975, p. 52-63) and Late Cretaceous time (Clark, 1977).

Evidence of worldwide oceanic anoxic events (Schlanger and Jenkyns, 1976) in the generally warm Mesozoic Era is indicated by remarkably widespread organic-rich black shales during the Late Jurassic and middle Cretaceous. The global extent and synchroneity of the Cretaceous anoxic events have been shown by the Deep Sea Drilling Project (Degens and Stoffers, 1976; Schlanger and Jenkyns, 1976; Ryan and Cita, 1977; Thiede and van Andel, 1977; Tissot et al, 1979). These paleoanoxic events have also been increasingly recognized on the continental shelves by petro­leum exploration drilling.

Identical worldwide oceanic anoxic events are inferred also to have prevailed intermittently in Paleozoic time because the widespread Lower Or-dovician, Silurian, and lower Carboniferous or­ganic-rich marine black shales coincide with times of widespread marine transgression and ice cap melting (Berry and Wilde, 1978).

Lastly, fiiere is statistical evidence that the known oil source bed systems present in the stratigraphic record are not randomly distributed in time but coincide with periods of worldwide transgression and oceanic anoxia (Arthur and Schlanger, 1979; Tissot, 1979). Prolific petroleum source beds deposited over part of the Jurassic and Cretaceous account for 70% of the oil in 148 investigated petroleum zones, which, themselves, total about 85% of the world's reserves (Tissot, 1979). Yet this privileged period for oil source bed deposition (Jurassic-Cretaceous) only amoimts to 17% of Phanerozoic time. Prolific petroleum re­gions where parts of the Cretaceous and/or Juras­sic play a major oil source role are the Middle East (Kent and Warman, 1972, p. 150), the North Sea (Ziegler, 1979), the West Siberian basin (Kon-torovich, 1971), and Mexico (Arthur and Schlan­ger, 1979).

CONCLUSIONS The explosive growth in knowledge by recent

observations in oceanography, geochemistry, and geomicrobiology has made it necessary to review many traditional concepts relating to oil source bed genesis. The proposed revised concepts fol­low.

1. Potential oil source beds are organic-rich sediments containing a kerogen type (type I or II) that is sufficiently hydrogen-rich to convert main­ly to oil during thermal maturation.

2. Measurement of organic carbon content, alone, is insufficient to identify potential oil source beds. For example, "gas-prone" transport­ed terrestrial organic matter (OM), oxidized aquatic OM deposited at high sedimentation rates, or reworked OM from a previous sedimen­tary cycle commonly create misleadingly high levels of organic carbon in ancient marine sedi­ments. High organic carbon content does not nec­essarily equate with oil source potential, hence precise measurement of kerogen type is indis­pensable to the identification of oil source beds. Insight into the early environmental factors that determine kerogen type, furthermore, is the prin­cipal key to understanding the temporal as well as paleogeographic distribution of oil source beds.

2. Organic matter in sediments below anoxic water (defined as water containing dissolved oxy­gen in amounts less than 0.5 ml/1) is commonly more reduced (hydrogen-rich), more lipid-rich, and more abundant than under oxic water.

3. The organic residues of anaerobic bacterial degradation are more reduced (hydrogen-rich) and more lipid-rich than those resulting from oxi­dation by aerobic microbes. Hence anaerobic res­idues are, potentially, the precursors to "oil-prone" kerogens. The latter are essential constitu­ents of potential oil source beds. The mechanisms leading to this preferential enrichment in lipids are little understood and in need of fundamental research in microbial biochemistry.

4. The overall range of organic carbon content under anoxic water is significantly wider and higher (about 1% to over 20% organic carbon) than that in sediments deposited under oxygenat­ed water. It is suggested that variations in sedi­mentation rate, if perhaps one of the underlying reasons for organic richness fluctuations under anoxic water, are not the main cause of enhanced preservation in reducing environments. Sedi­ments deposited under anoxic marine wato- also contain abnormally high concentrations of U, Cu, Mo, Ni, P, and S which tend to correlate posi­tively with organic carbon concentrations. Organ­ic matter deposited under anoxic water, wherever investigated, is hydrogen-rich and therefore "oil-prone."

5. Organic carbon content in sediments under oxygenated water is primarily a function of sedi­mentation rates up to the point (yet ill-defined) where organic matter dilution significantly de­presses concentrations. Surface primary produc­tivity appears to have little influence on organic carbon concentrations in sediments under "oxic" water because it is systematically compensated by organic matter consumption at the seafloor by benthic organisms. Organic carbon contents un-

Anoxic Environments 1205

der "oxic" water typically range between less than 0.5% to maxima between 3 and 4%, regard­less of surface productivities. Organic matter de­posited under oxic water, wherever investigated, is hydrogen depleted and therefore "gas prone."

6. Bioturbation, or biologic sediment mixing by metazoans, is ubiquitous at all water depths un­der "oxic" water. It causes acceleration of geo-chemical processes at and below the sediment-water interface. Added to this effect is the grazing of the sediment bacterial biomass by the mobile deposit feeders (burrowers) responsible for bio­turbation. Bioturbated sediments under "oxic" water are homogenized whereas sediments depos­ited under anoxic water are varved or laminated and devoid of burrowing, deposit-feeding infau-na.

7. Maximum oxygen saturation in seawater is about 6 to 8.5 ml/1, but benthic metazoans are unaffected by lower oxygen concentrations, down to about 1 ml/1 of water. Between 0.1 and 0.5 ml/l, a low diversity, highly stressed, suspension-feeding (non-burrowing) epifauna can still sur­vive above the sediment-water interface. Below 0.5 to 0.3 ml/1, bioturbation by deposit feeders (burrowers) virtually ceases and the overall benthic biomass is sharply reduced. Below 0.1 ml/1, all metazoans disappear leaving anaerobic bacteria as the only effective reworkers. Hence 0.5 ml/1 (the threshold of arrested bioturbation), not total absence of oxygen in water, is proposed as the effective "biochemical fence" between po­tentially good or poor quahtative and quantita­tive organic matter preservation.

8. The tolerance of protozoans, such as benthic foraminifers, to very low oxygen concentrations appears to be higher than that of metazoans. The ecology of benthic foraminifers under oxygen-de­pleted water is clearly in need of fundamental re­search.

9. Anoxic water is defined, for this study, as water containing less than 0.5 miUiliters oxygen per hter of water. Anoxic conditions occur where the natural demand for oxygen in water exceeds the supply. Oxygen demand relates to surface pri­mary productivity, whereas oxygen supply below the depth of surface mixing depends on water cir­culation.

10. Lakes are rarely permanently anoxic at depth. It takes special conditions such as overfer-tilization, lack of overturning of the water in warm climates, and preferably, but not necessar­ily, deep water to create stable anoxic conditions. When they occur, as in Lake Tanganyika (east Africa), sediments rich in organic matter are de­posited. Warm paleoclimatic conditions with lack of seasonal contrast and moderate rainfall should

be most favorable to lacustrine oil source bed de­position.

11. At present, the majority of silled basins, whether landlocked or on open shelves, are not anoxic and thus do not favor deposition of oil source bed precursor sediments. Only those silled basins that have a water circulation pattern im­plying a "positive water balance" become anoxic, for example, the Black Sea or the Baltic Sea. The degree of ventilation in silled depressions on open oceanic shelves follows that one prevailing at comparable depth in the adjoining open oceanic water masses. Hence silled-basin geometry in an­cient sedimentary basins does not automatically imply the presence of oil source beds. Silled ba­sins should be prone to anoxia and oil source bed deposition at times of worldwide transgression, both at high and low paleolatitudes. At times of worldwide regression, those landlocked basins in temperate paleoclimatic belts with high rainfall should be more favorable than those in arid zones.

12. Coastal upweUings, which are commonly associated with high primary productivity of the surface layers, do not originate from great water depths. Only upwellings associated with anoxic layers show any abnormal organic preservation as well as phosphorite deposition in bottom sedi­ments (e.g., Benguela and Peru Currents). Exten­sive oil source beds and phosphorites resulting from upwelling conditions should be sought at low paleolatitudes and preferably at times of worldwide transgression.

13. Large volumes of anoxic water are present in the oxygen-minimum layers of open oceans such as the Indian Ocean without silling or physi­cal barriers. Significant organic enrichment in sediments has been found wherever the oxygen-minimum layer reaches the anoxic level and inter­sects the continental slope and shelf. The prefer­ential distribution of oxygen-minimum layers in the world ocean is at low latitudes and on the western side of continents, primarily because of deep oceanic circulation spurred and controlled by climatic contrast of the poles and the equator and the Coriolis force. Inferred are similar global patterns, determining the position of oxygen-min­imum layers during the geologic past, but with much more widespread extensions of anoxia in­vading whole oceanic basins. This is suggested by remarkably widespread organic-rich marine black shales at times of global climatic warm-ups and major transgressions as in Late Jurassic and mid­dle Cretaceous time. Known marine oil source bed systems in the stratigraphic record are not randomly distributed in time but coincide with periods of worldwide transgression and oceanic

1206 G. J. Demaison and G. T. Moore

ANOXIC BASIN TYPE

ANOXIC LAKES

ANOXIC SILLED BASINS

ANOXIC LAYERS W/ UPWELLINGS

ANOXIC OPEN OCEAN

PALE0GE06RAPHIC SETTING

EQUABLE. WARM. RAINY. EARLY RIFTS. INTERMOUNTAIN BASINS.

TEMPERATE TO WARM. RAINY. INTRACRATONIC SEAS. ALSO ANOXIC POCKETS ON SHELVES.

OCEANIC SHELVES AT LOW LATITUDES. WEST SIDE OF CONTINENTS.

BEST DEVELOPED AT TIMES OF GLOBAL WARM-UPS & MAJOR TRANSGRESSIONS.

STRATI6RAPHIC DISTRIBUTION OF ANOXIC SEDIMENTS

CONTINUOUS WITHIN SAME ANOXIC LAKE SYSTEM.

VARIABLE. TENDS TO BE RICHEST AT BOTTOM OF BASIN OR POCKET.

OFTEN NARROW TRENDS. BUT CAN BE WIDESPREAD. PHOSPHORITES. DIATOMITES.

VERY WIDESPREAD W/LITTLE VARIATION. OFTEN SYNCHRONOUS WORLDWIDE.

FIG. 18—Classification of anoxic depositional models. Certain settings in marine realm may combine two or even three anoxic basin types, for example, local anoxic layers caused by upwelling may reinforce anoxic open-ocean oxygen-minimum layer impinging upon shelf with silled depressions.

anoxia. 14. Geochemistry now provides the means to

identify paleoanoxic events in the stratigraphic record. Paleogeographic recognition of the pro­posed anoxic models—large anoxic lakes, anoxic silled basins, anoxic layers caused by upweUing, and anoxic open ocean—in ancient sedimentary basins should help geologists and geochemists in the regional mapping of oil shales and oil source beds as shown in Figure 18.

REFERENCES CITED

Aller, R. C, 1978, The effects of animal-sediment inter­actions on geochemical processes near the sediment-water interface, in Estuarine interactions: New York, Academic Press, p. 157-172.

Arthur, M. A., and S. O. Schlanger, 1979, Cretaceous "oceanic anoxic events" as causal factors in devel­opment of reef-reservoired giant oil fields: AAPG BuU., V. 63, p. 870-885.

BeUaeva, A. N., and E. A. Romankevich, 1976, Chemi­cal conversion of lipids during oceanic sedimentogen-esis, in N. B. Vassoevich, ed., Issledovania Organish-eskogo Veshchestva Sovremenyk Iskopaemyk Osadkov: Moscow, Adka. Nauk SSSR, p. 81-102 (in Russian).

Berggren, W. A., and C. D. HoUister, 1977, Plate tec­tonics and paleo-circulations—commotion in the ocean: Tectonophysics, v. 38, p. 11-4S.

Berry, W. B. N., and P. Wilde, 1978, Progressive venti­lation of the oceans: Am. Jour. Sci., v. 278, p. 257-275.

Bezrukov, P. L., et al, 1977, Organic carbon in the upper sedimentary layer of the world ocean: Oceanology, v. 17, no. 5, p. 561-564.

Bordovsky, O. K., 1%5, Accumulation and transforma­tion of organic substance in marine sediments: Ma­

rine Geology, v. 3, p. 3-114. Brongersma-Sanders, M., 1966, The fertiUty of the sea

and its bearing on the origin of oil: Advancement of Science, v. 23, p. 41-46.

1972, Hydrological conditions leading to the de­velopment of bituminous sediments in the pre-evapo-rite phase, in Geology of saline deposits: Unesco Earth Sci. Ser., no. 7, p. 19-21.

Bubnov, V. A., 1966, The distribution pattern of mini­mum oxygen concentrations in the Atlantic: Ocean­ology, V. 6, p. 193-201.

Byers, C. W., and D. W. Larson, 1979, Paleoenviron-ments of Mowry Shale (Lower Cretaceous), western and central Wyoming: AAPG Bull., v. 63, p. 359-361.

Calvert, S. E., 1964, Factors affecting distribution of laminated diatomaceous sediments in Gulf of CaU-fomia, in Marine geology of the Gulf of CaUfomia: AAPGMem. 3, p. 311-330.

1976, The mineralogy and geochemistry of near-shore sediments, in J. P. Riley et al, eds.. Chemical oceanography, v. 6, 2d ed.: New York, Academic Press, p. 187-280.

' and N. B. Price, 1971a, Upwelling and nutrient regeneration in the Benguela Current, October, 1968: Deep-Sea Research, v. 13, p. 505-523.

1971b, Recent sediments of the South­west African Shelf, in Atlantic continental margins: London, Inst. Geol. Sci., p. 175-185.

Clark, D. L., 1977, Climatic factors of the late Mesozoic and Cenozoic Arctic Ocean, in Polar Oceans: Cal­gary, Alberta, Arctic Inst. North America, p. 603-615.

Claypool, G. E., and I. R. Kaplan, 1974, The origin and distribution of methane in marine sediments, in I. R. Kaplan, ed.. Natural gases in marine sediments: New York, Plenum Press, p. 99-140.

A. H. Love, and E. K. Maughan, 1978, Organic geochemistry, incipient metamorphism, and oil gener­ation in black shale members of Phosphoria Forma-

Anoxic Environments 1207

tion. Western Interior United States: AAPG Bull., v. 62, p. 98-120.

Clementz, D. M., G. J. Demaison, and A. R. Daly, 1979, Well site geochemistry by programmed pyroly-sis: Offshore Technology Conf., v. 1, p. 465-470.

Combaz, A., and R. Pelet, eds., 1978, Geochimie orga-nique des sediments marins profonds, ORGON II, Atlantique-N.E. Bresil: Paris, Editions CNRS, 456 p.

E>ebyser, Y., R. Pelet, and M. Dastillung, 1977, Geochi­mie organique des sediments marins recents, Mer Noire, Baltique, Atlantique (Mauritanie), in Ad­vances in organic geochemistry, 1975: Madrid, ENA-DIMSA, p. 289-320.

Degens, E. T., 1974, Cellular processes in Black Sea sed­iments, in The Black Sea: AAPG Mem. 20, p. 296-307.

and K. Mopper, 1976, Factors controlling the distribution and early diagenesis of organic material in marine sediments, in J. P. Riley et al, eds.. Chemi­cal oceanography, v. 6, 2d ed.: New York, Academic Press, p. 59-113.

and D. A. Ross, eds., 1974, The Black S e a -geology, chemistry and biology: AAPG Mem. 20, 633 P-

and P. Stoffers, 1976, Stratified waters as a key to the past: Nature, v. 263, p. 22-27.

R. P. Von Herzen, and H.-K. Wong, 1971, Lake Tanganyika; water chemistry, sediments, geological structure: Naturwissenschaften, v. 58, p. 224-291.

- et al, 1973, Lake Kivu; structure, chemistry and biology of an East African rift lake: Geol. Rund­schau, v. 62, p. 245-277.

- et al, 1978, Varve chronology: estimated rates of sedimentation in the Black Sea deep basin: Initial Rept. Deep Sea Drilling Project, v. 42, pt. 2, p. 499-508.

Demaison, G. J., and M. Shibaoka, 1975, Contribution to the study of hydrocarbon generation from hydro­gen-poor kerogens: 9th World Petroleum Cong., To­kyo, Proc., V. 2, p. 195-197.

Denser, W. G., 1971, Organic-carbon budget of the Black Sea: Deep-Sea Research, v. 18, p. 995-1004.

1975, Reducing environments, in J. P. Riley and R. Chester, eds.. Chemical oceanography, v. 3: New York, Academic Press, p. 1-37.

Dickert, P. F., 1966, Neogene phosphatic facies in Cali­fornia: PhD thesis, Stanford Univ.

Didyk, B. M., et al, 1978, Organic geochemical indica­tors of paleoenvironmental conditions of sedimenta­tion: Nature, v. 272, p. 216-222.

Dow, W. G., 1977, Kerogen studies and geological in­terpretations: Jour. Geochem. Exploration, v. 7, p. 79-99.

and D. B. Pearson, 1975, Organic matter in Gulf Coast sediments: 7th Offshore Technology Conf., Proc., v. 3, OTC 2343, p. 85-94.

Emery, K. O., and E. Uchupi, 1972, Western North At­lantic Ocean: topography, rocks, structures, water, life, and sediments: AAPG Mem. 17, 532 p.

Espitalie, J., et al, 1977, Source rock characterization method for petroleum exploration: Offshore Technol­ogy Conf. Proc., v. 3, OTC 2935, p. 439-444.

Fairbridge, R. W., ed., 1966, The encyclopedia of oceanography: New York, Van Nostrand-Reinhold,

1021 p. Foree, E. G., and P. L. McCarty, 1970, Anaerobic de­

composition of algae: Environ. Sci. Technology, v. 4, p. 842-849.

Gaskell, S. J., et al, 1975, The geochemistry of a recent marine sediment off northwest Africa. An assessment of source of input and early diagenesis: Deep-Sea Re­search, V. 22, p. 777-789.

Gerschanovich, D. E., V. V. Veber, and A. E. Konjuk-hov, 1976, Organic matter in the bottom sediments of the Peru region of the Pacific Ocean, in N. B. Vassoe-vich, ed., Issledovania Organisheskogo Sovremenyk Iskopaemyk Osadkov: Moscow, Akad. Nauk, SSSR, p. 121-128 (in Russian).

Gormly, J. R., and W. M. Sackett, 1977, Carbon isotope evidence for the maturation of marine lipids, in Ad­vances in organic geochemistry, 1975: Madrid, ENA-DIMSA, p. 321-339.

Gorshkov, S. G., 1974, World ocean atlas, v. 1, Pacific Ocean: New York, Pergamon Press, 338 p. (in Rus­sian).

1979, World ocean atlas, v. 2, Atlantic and Indi­an Oceans: New York Pergamon Press, 350 p. (in Russian).

Grasshoff, K., 1975, The hydrochemistry of landlocked basins and fjords, in J. P. Riley and J. Skirrow, eds.. Chemical oceanography, v. 2, 2d ed.: New York, Ac­ademic Press, p. 456-597.

Gross, G. M., et al, 1972, Distribution of organic carbon in surface sediment, northeast Pacific Ocean, in The Columbia River Estuary and adjacent ocean waters: Seattle, Univ. Washington Press, p. 254-264.

Hallam, A., 1975, Jurassic environments: Cambridge, U.K., Cambridge Univ. Press, 269 p.

Hartmann, M., et al, 1971: Oberflachensedimente im Persischen Gulf und Gulf von Oman: "Meteor" For-schungsergebnisse, Reihe C, No. 4, 76 p.

et al, 1976, Chemistry of late Quaternary sedi­ments and their interstitial waters from the NW Afri­can continental margin: "Meteor" Forschungaser-gebnisse, Reihe C, No. 24, p. 1-67.

Heath, G. R., T. C. Moore, Jr., and J. P. Dauphin, 1977, Organic carbon in deep-sea sediments, in R. N. An­dersen, ed., The fate of fossil fuel CO2 in the oceans: New York, Plenum Press, p. 627-639.

Hirst, D. M., 1974, Geochemistry of sediments from eleven Black Sea cores, in The Black Sea: AAPG Mem. 20, p. 430455.

HoUister, C. D., R. Flood, and I. W. McCave, 1978, Geological aspects of the benthic boundary layer: Oceanus, v. 21, p. 5-13.

Hue, A-Y., 1978, Geochimie organique des schistes bi-tumineux du Toarcian du Bassin de Paris: PhD the­sis, Univ. Louis Pasteur, Strasbourg, France.

Hunt, J. M., 1979, Petroleum geochemistry and geol­ogy: San Francisco, W. H. Freeman and Co., 498 p.

Idyll, C. P., 1973, The anchovy crisis, reprinted 1977, in Ocean science; readings from Scientific American: San Francisco, W. H. Freeman and Co., p. 223-230.

Jenkins, O. P., ed., 1943, Geologic formations and eco­nomic development of the oil and gas fields of CaU-fomia: California Div. Mines Bull. 118, 778 p.

Jones, G. E., 1973, An ecological survey of open ocean and estuarine microbial populations. I, TTie impor-

1208 G. J. Demaison and G. T. Moore

tance of trace metal ions to microorganisms in the sea, in L. H. Stevenson and R. R. Colwell, eds., Es-tuarine microbial ecology: Columbia, S.C, Univ. South Carolina Press, p. 233-241.

Kent, P. E., and H. R. Warman, 1972, An environmen­tal review of the world's richest oil-bearing region; the Middle East: 24th Intemat. Geol. Cong., Montre­al, Proc., sec. 5, p. 142-152.

Klingebiel, A., 1976, Sediments et milieuz sedimentaires dans le Golfe du Benin: Centre Recherche Pau-SNPA, Bull. 10, p. 129-148.

Koblenz-Mishke, O. I., V. V. Volkonsky, and J. G. Ka-banova, 1970, Plankton primary production of the world ocean, in W. S. Wooster, ed.. Symposium on scientific exploration of the southern Pacific, Scripps Inst Oceanog., 1968, 9th Mtg. SCOR: Washington, D.C., Natl. Acad. Sci., p. 183-193.

Konjukhov, A. E., 1976, Fades characteristics of con­temporary sediments from the western submarine margin of the Indian Peninsula, in N. B. Vassoevich, ed., Issledovania Organisheskogo Veshchestva Sovre-menyk Iskopaemyk Osadkov: Moscow, Akad. Nauk SSSR, p. 111-120 (in Russian).

Kontorovich, A. E., ed., 1971, Geochemistry of petro­leum bearing Mesozoic formations in the Siberian ba­sins: USSR Ministry of Geology Publication, SNIIG-IMS Bull. 118, 85 p. (in Russian).

Krumbein, W. C , and R. M. Garrels, 1952, Origin and classification of chemical sediments in terms of pH and oxidation-reduction potentials: Jour. Geology, v. 60, p. 1-33.

Lijmbach, G. W. M., 1975, On the origin of petroleum: 9th World Petroleum Cong., Tokyo, Proc, v. 2, p. 357-368.

Lisitzin, A. P., 1972, Sedimentation in the world ocean: SEPM Spec. Pub. 17, 218 p.

Logvinenko, N. V., and Ye A. Romankevich, 1973, Re­cent sediments of the Pacific Ocean on the coast of Peru and Chile: Lithology and Mineral Resources, v. 8, p. 1-11; also in Litol. Polez. Iskop., no. 1, p. 3-16, 1973 (in Russian).

Manheim, F. T., 1976, Interstitial waters of marine sedi­ments, in J. P. Riley et al, eds., Chemical oceanogra­phy, V. 6, 2d ed.: New York, Academic Press, p. 115-186.

Momper, J. A., and J. A. Williams, 1979, Geochemical exploration in Powder River basin, northeastern Wy­oming and southeastern Montana (abs.): AAPG Bull., V. 63, p. 497.

Morris, K. A., 1979, A classification of Jurassic marine shale sequences: an example from the Toarcian (Lower Jurassic) of Great Britain: Paleogeography, Paleoclimatology, Paleoecology, v. 26, p. 117-120.

Nissenbaum, A., M. J. Baedecker, and I. R. Kaplan, 1972, Dissolved organic matter from interstitial or­ganic water of a reducing marine fjord, in Advances in organic geochemistry, 1971: Intemat. Ser. Mon. Earth Sci., v. 33, p. 427-440.

Nixon, R. P., 1973, Oil source beds in Cretaceous Mow-ry Shale of Northwestern Interior United States: AAPG BuU., V. 57, p. 136-161.

Oremland, R. S., and B. F. Taylor, 1978, Sulfate reduc­tion and methanogenesis in marine sediments: Geo-chim. et Cosmochim. Acta, v. 42, p. 209-214.

Orr, W. L., and A. G. Gaines, 1974, Observations on rate of sulfate reduction and organic matter oxidation in the bottom waters of an estuarine basin: the upper basin of the Pettaquamscutt River (Rhode Island), in Advances in organic geochemistry, 1974: Paris, Tech-nip, p. 790-812.

Palfrey, M. K., and G. G. Day, 1968, Oceanography of Baffin Bay and Nares Strait in the summer of 1968: Washington, D.C., Coast Guard Rept. 373-16.

Parker, R. H., 1964, Zoogeography and ecology of ma­cro-invertebrates of GtUf of California and continen­tal slope of westem Mexico, in Marine geology of the Gulf of California: AAPG Mem. 3, p. 331-376.

Parrish, J. T., K. S. Hansen, and A. M. Ziegler, 1979, Atmospheric circulation and upwelling in Paleozoic, with reference to petroleum source beds (abs.): AAPG Bull., V. 63, p. 507-508.

Pelet, R., and Y. Debyser, 1977, Organic geochemistry of Black Sea cores: Geochim. et Cosmochim. Acta, v. 41, p. 1575-1586.

Peng, T. H., et al, 1977, Benthic mixing in deep-sea cores as determined by C'* dating and its implication regarding climate, stratigraphy and the fate of fossil fuel CO2, in N. R. Andersen, ed.. The fate of fossil fuel CO2 in the oceans: New York, Plenum Press, p. 355-373.

Phleger, F. B, 1964, Patterns of living benthonic Foram-inifera, Gulf of California, in Marine geology of the Gulf of California: AAPG Mem. 3, p. 377-394.

Redfield, A. C , 1958, Preludes to the entrapment of organic matter in the sediments of Lake Maracaibo, in L. G. Weeks, ed., Habitat of Oil: AAPG, p. 968-981.

Reid, J. L., 1965, Intermediate waters of the Pacific Ocean: Baltimore, Johns Hopkins Univ. Press, 85 p.

Rhoads, D. C , 1974, Organism-sediment relationship on the muddy sea floor: Oceanogr. Marine Biology Ann. Rev., v. 12, p. 263-300.

Richards, F. A., 1970, The enhanced preservation of organic matter in anoxic marine environments, in Symposium on organic matter in natural waters: Univ. Alaska Inst. Marine Sci. Occasional Pub. 1, p. 399-411.

1976, The Cariaco basin (trench): Oceanogr. Marine Biology Ann. Rev., v. 13, p. 11-67.

and R. F. Vaccaro, 1956, The Cariaco Trench, an anaerobic basin in the Caribbean Sea: Deep-Sea Res., V. 7, p. 163-172.

Roden, G. I., 1964, Oceanographic aspects of Gulf of California, in Marine geology of the Gulf of Califor­nia: AAPG Mem. 3, p. 30-58.

Roehler, H. W., 1974, Depositional environment of rocks in the Piceance Creek basin, Colorado: Rocky Mtn. Assoc. Geologists Field Conf. Guidebook 25, p. 57-64.

Romankevich, E. A., 1977, Organic geochemistry of oceanic sediments: Moscow, Izd. Nauka, p. 13 (in Russian).

Rosenberg, R., 1977, Benthic macrofaunal dynamics, production and dispersion on an oxygen-deficient es­tuary of west Sweden: Jour. Exp. Marine Biology Ec­ology, v. 26, p. 107-133.

Ross, D. A., and E. T. Degens, 1974, Recent sediments of Black Sea, in The Black Sea: AAPG Mem. 20, p.

Anoxic Environments 1209

78. Theede, H., et al, 1969, Studies on the resistance of ma­

rine bottom invertebrates to oxygen deficiency and hydrogen sulfide: Mar. Biology, v. 2, p. 325-337.

Thiede, J., and T. H. van Andel, 1977, The paleoenvi-ronment of anaerobic sediments in the late Mesozoic South Atlantic Ocean: Earth and Planetary Sci. Let­ters, V. 33, p. 301-309.

Tissot, B. P., 1979, Effects on prolific petroleum source rocks and major coal deposits caused by sea-level changes: Nature, v. 277, p. 377-380.

and D. H. Welte, 1978, Petroleum formation and occurrence: New York, Springer-Verlag, 521 p.

G. Deroo, and A. Hood, 1978, Geochemical study of the Uinta basin; formation of petroleum from the Green River Formation: Geochim. et Cos-mochim. Acta, v. 42, p. 1469^1486.

- et al, 1974, Influence of nature and diagenesis of organic matter in formation of petroleum: AAPG BuU., V. 58, p. 499-506.

et al, 1979, Paleoenvironment and petroleum potential of mid-Cretaceous black shales in Atlantic basins (abs.): AAPG Bull., v. 63, p. 542.

Turekian, K. K., J. K. Cochran, and D. J. De Master, 1978, Bioturbation in deep-sea deposits: rates and consequences: Oceanus, v, 21, p. 34-41.

Tyson, R. V., R. C. L. Wilson, and C. Downie, 1979, A stratified water column enviromnental model for the type Kimmeridge clay: Nature, v. 277, p. 377-380.

Udintsev, Gb., ed., 1975, Geological-geophysical atlas of the Indian Ocean: New York, Pergamon Press, 151 P-

Usher, J. L., and P. Supko, 1978, Initial reports of the Deep Sea Drilling Project, v. 42, pt. 2, 1243 p.

Vail, P. R., and R. M. Mitchum, 1979, Global cycles of sea level change and their role in exploration: 10th World Petroleum Cong., Preprint, Bucharest Panel Discussion 2, Paper 4.

van Andel, T. H., 1964, Recent marine sediments of Gulf of California, in Marine geology of Gulf of Cali­fornia: AAPG Mem. 3, p. 216-310.

Veber, V. V., 1977, Generation of hydrocarbon gases in modem sedimentary formation of the eastern and southern parts of the Atlantic Ocean: Geologyia Nef-ti i Gaza, no. 1, p. 59-64.

Veeh, H. H., 1979, Modem environment of phosphorite formation and the geochemical balance of phospho­rus in the ocean, in P. J. Cook and J. H. Shergold, eds., Proterozoic-Cambrian phosphorites: Project 156 of UNESCO-IUGS: Canberra, Australian Natl. Univ. Press, p. 59.

W. C. Bumett, and A. Soutar, 1973, Contempo­rary phosphorites on the continental margin of Pern: Science, v. 181, p. 844-845.

• S. E. Calvert, and N. B. Price, 1974, Accumula­tion of uranium in sediments and phosphorites on the southwest African shelf: Marine Chemistry, v. 2, p. 189-202.

Wolfe, R. S., 1971, Microbial formation of methane: Adv. Microbial Physiology, v. 6, p. 107-146.

Wyrtki, K., 1962, The oxygen minimum in relation to ocean circulations: Deep-Sea Research, v. 9, p. 11-23.

Ziegler, P. A., 1979, Factors controlling North Sea hy­drocarbon accumulations: World Oil, v. 189, no. 6, p. 111-124.

J. T. Parrish, and R. C. Humphreville, 1979, Pa-leogeography, upwelling and phosphorites (abs.), in P. J. Cook and J. H. Shergold, eds., Proterozoic-Cam­brian phosphorites. Project 156 of UNESCO-IUGS: Canberra, Australian Natl. Univ. Press, p. 21.

183-199. Ryan, W. B. F., and M. B. Cita, 1977, Ignorance con­

cerning episodes of ocean-wide stagnation: Marine Geology, v. 23, p. 197-215.

Sackett, W. M., et al, 1978, A carbon inventory for Orca brines and sediments: Earth and Planetary Sci. Let­ters, V. 44, p. 73-81.

SaUot, A., 1977, Etude comparative des acides gras de I'eau et de I'eau interstitieUe a Tinterface eau de mer profonde/sediment: Acad. Sci. Comptes Rendus, ser. D-857, p. 285 (10 Oct. 1977).

Schaefer, W., 1972, Ecology and palaeoecology of ma­rine environments: Chicago, Univ. Chicago Press, 568 p.

Schlanger, S. O., and H. C. Jenkyns, 1976, Cretaceous oceanic anoxic events: causes and consequences: Ge­ologic en Mijnbouw, v. 55, p. 179-184.

Scranton, M. I., and J. W. Farrington, 1977, Methane production in the waters off Widvis Bay: Jour. Geo-phys. Research, v. 82, p. 4947-4953.

Shimkus, K. M., and E. S. Trimonis, 1974, Modem sedi­mentation in Black Sea, in The Black Sea: AAPG Mem. 20, p. 249-278.

Shokes, R. F., et al, 1977, Anoxic hypersaline basin in the northem Gulf of Mexico: Science, v. 196, p. 1443-1446.

Smith, J. W., and W. A. Robb, 1973, Aragonite and the genesis of carbonates in Mahogany zone oil shales of Colorado's Green River Formation: U.S. Bur. Mines Rept. 7727, 21 p.

Spencer, D. W., H. Susumu, and D. G. Brewer, 1978, Particles and particle fluxes in the ocean: Oceanus, v. 21, p. 20-25.

Stackelberg, U. V, 1972, Faziesverteilung in Sedimenten des indisch-pakistanischen Kontinentalrandes (Ara-bisches Meer): "Meteor" Forschungsergebnisse, Rei-he C, No. 9, p. 1-173.

Stanley, D. J., 1978, Ionian Sea sapropel distribution and late Quaternary paleo-oceanography in the east-em Mediterranean: Nature, v. 274, p. 149-151.

Stevenson, F. J., and C. N. Cheng, 1972, Organic geo­chemistry of the Argentine basin sediments: carbon-nitrogen relationships and Quaternary correlations: Geochim. et Cosmochim. Acta, v. 36, p. 653-671.

Stommel, H., 1958, The circulation of the abyss, reprint­ed, 1977, in Ocean science, readings from Scientific American: San Francisco, W. H. Freeman and Co., p. 139-144.

Summerhayes, C. P., U. de Melo, and H. T. Baretto, 1976, The influence of upwelling on suspended mat­ter and shelf sediments off southeastern Brazil: Jour. Sed. Petrology, v. 46, p. 819-828.

Swain, F. M., 1970, Non-marine organic geochemistry: Cambridge, U.K., Cambridge Univ. Press, 445 p.

Swanson, V. E., 1960, Oil yield and uranium content of black shales: U.S. Geol. Survey Prof. Paper 356-A, p. 1-44.

TaUing, J. F., 1963, Origin of stratification in an African rift lake: Limnology and Oceanography, v. 8, p. 68-